TRACERS IN THE SEA
w. S. Broecker and T.-H. PengTRACERSIN THE SEAWALLACE SMITH BROECKER
COLUMBIA UNIVERSI TY

AND
TSUNG-HUNG PENGOAK RIDGE NATIONAL LABORATORYA PUBLICATION OF THE LAMONT-DOHERTY GEOLOGICAL OBSERVATORY,
COLUMBIA UNIVERSITY, PALISADES, NEW YORK 10964

Printed in the United States of America

Tsung-Hung Peng's work at Oak Ridge was sponsored jointly by the
National Science Foundation's Ecosystem Studie~ Program under Interagency Agreement No. DEB 8115316 and the Carbon Dioxide Research D1vision, Office of Energy Research, U.S. Department of Energy, under contract W-740S-eng-26 with Union Carbide Corporation.

A decade has passed since Chemical Oceanography, the predecessor to this book, was written.
During this time our knowledgeof the chemistry of the oceans has mushroomed.
The Geochemical
Ocean Sections program (GEOSECS) fleshed out our sketchy pictureof the distributions of chemical, isotopic and radiochemical
tracers in the sea.
The discovery of ridge crest hydrothermal
systems has added a new dimension to ocean chemistry. The resultsof sediment trap deployments, of pore water measurements, and ofin situ filtrations of particulate matter have expanded our knowledge of the cycle of particulate matter in the sea. Hence, rather
than revising Chemical Oceanography, we decided to write a newbook, Tracers in the Sea. While this new book follows the outlineof its predecessor and contains material from its predecessor,
about 75% of the material is new.
Unlike Chemical Oceanography,
which could be read from cover to cover in a few sittings, Tracers
in the Sea has far more meat and cannot be so quickly digested.
The objective of this book is to demonstrate how the distribution of tracers in the sea and its sediments can be used to aid
us in deciphering the operations plan and history of operation forthe ocean as a chemical plant.
To do this requires the integration of chemical information with information based on studies ofocean sediments, of organisms living in the sea, and of ocean currents.
This integration leads to a description of how the ocean"works" from the point of view of a geochemist. Such integrations
are not common to the field of oceanography because only in attempting to understand the chemistry of the sea is it necessary toconsider the results from all the sister disciplines.
One dominant theme which recurs in the book is the operationof the earth's carbon cycle. Carbon is a key element in organisms
and their shells.
It is a constituent of the key greenhouse gas
in our earth's atmosphere.
It has two isotopes of great importance as tracers, namely, 14C and 13 C • In the book we explore not
only the present day distribution of this key element but also how
this distribution may have differed during the last glacial period
and how it will differ during the coming fossil fuel C02-induced
superinterglacial period.
While aimed both at the professional and at the student, thebook has the format of a text.
As such it puts more emphasis on
explanation than on documentation.
To counter the somewhat thinreferencing in the body of the text a rather elaborate reference
list is given at the end of the book.
This list contains over
700 entries and is organized by topic.
Thus, a reader interested
in finding all the papers which bear on, let's say, the distribu-

tion of 3 He or of lead in the sea has only to look up the sections
on these properties in the reference list to find a fairly complete list of papers written on these subjects.
As the papers oneach topic are listed in order of publication date and as they are
annotated with a sentence or two stating their contribution, thereader can get some sense of development of the data base andthinking in a given area by looking through this reference lis t.
Hence, the reference numbers in the text and in the figure andtable captions serve a dual purpose.
They give the reader thesource of specific information used in the book and they lead him
to the section in the references which lists related papers not
specifically mentioned in the text.
Every text book in science needs problems.
They reinforce
the concepts in the text and force the student to think carefully
about the quantitative aspects of the sUbject.
We include boththe usual short problems and also what we call superproblems. The
latter give the student a better in-depth view of key material
and also good training in box modeling which plays such animportant role in marine geochemistry.
As the superproblems
require 4-10 hours of work, the instructor should consider haVing
students work in pairs on the problems and having each student
work on only 1 or 2 of the problems during a semester. We havefound that presentations of the solutions by the students doing
the work to the rest of the class proves very effective.
Solutions to both the problems and superproblems can be obtainedby writing to the authors.
Much of this book was written during a sabbatical leave inGermany.
I (WSB) would like to thank the Alexander Von Humboldt
Foundation for providing fellowship support for a nine month visit
to Heidelberg and to Karl Otto Munnich and Wolfgang Roether andtheir colleagues in the Umweltphysik group for their friendshipand help during this time.
The drafting for the book was done by
Patty Catanzaro and the typing by Vicky Costello, Rae Pochapsky
and Andrea Radigan.
Ellen C:oxe acted as liaison between theseproduction efforts at Lamont and the writing and modeling activities by T.-H. Peng at Oak Ridge National Laboratory and by
W.S. Broecker in Heidelberg.
The sketch of Arnold Bainbridge andthe cover design were done by Dee Breger, a member of thetechnical staff in geochemistry at Lamont-Doherty.
The coverphoto,arosette
cast
from
the
Woods
Hole
Oceanographic
Institution research vessel KNORR, was taken by David Chipman of
the Lamont-Doherty Geochemistry group.
We thank Heinrich Holland for encouraging us to follow arather unconventional plan for publication.
The Lamont-Doherty
Geological Observatory served as publisher. Author-ready copy was
prepared using the geochemistry group word processor under thesupervision of Patty Catanzaro.
The printing was jobbed out to
Edwards Brothers Incorporated, Ann Arbor, Michigan. The advertising and distribution were handled by Observatory staff.
We thank
IT. Barry Raleigh, director of the Observatory, for his courageousbacking of this dangerous experiment.

CHAPTER 1 INTERNAL CYCLING AND THROUGHPUT
Pathways from River Mouth to Sea Floor
Introduction

1

Depth Profiles of Sea Salt Composition

2

Constituent Classification

6

The Chemical Composition of Marine Organic Matter

8

Composition of Particulate Matter Caught in Sediment Traps

13

A Simple Model for Biologically Utilized Constituents15

The Distributions of Biointermediate CDnstituents

22

Estimation of Input Rates

25

Horizontal Segregation of Constituents in the Deep Sea

28

Summary40

CHAPTER 2 THE SEDIMENTARY SINK
Factors Influencing the Distribution of Sedimentary ConstituentsIntroduction

45

Sediment Types

45

Distribution of Opal Production

47

Opal Solution on the Sea Floor

49

Distribution of Qalcite in Marine Sediments

58

Degree of Calcite Saturation

61

Variation in the Carbonate Ion Qontent of Sea Water

62

Spacial Variations in the CaC0 3 Saturation of Sea Water

71

Factors Controlling the Rate of Calcite Solution

84

Thickness and Shape of the SUblysocline Transition Zone

89

Variation of Sediment Type with Time94

Manganese Nodules

98

Summary106
CHAPTER 3 THE ATMOSPHERIC IMPRINT
The Cycles of Gases within the Sea

Introduction

110

Solubilities of Gases in Sea Water

111

The Rate of Gas Exchange

113

Stagnant Film Thickness Derived from Natural Radiocarbon

118

Stagnant Film Thicknesses Determined by the Radon Method

122

Oxygen Concentrations in Surface Ocean Water

127

Oxygen Deficiencies in the Deep Sea

131

The Marine NzO Cycle

139

Excess Helium

147

The Carbon Dioxide Content of Surface Ocean Water

149

Origin of the Equatorial Pacific COz Anomaly

158

Summary

161

CHAPTER 4 REACTIVE METALS AND THE GREAT PARTICULATE SWEEPThe Cycle of Metals in the Sea
Introduction

166

Products of Uranium and Thorium Decay

168

Thorium Isotopes in the Sea

172

Protactinium-231 to Thorium-230 Activity Ratios

184

The Distribution of Lead-210

189

The Distribution of Polonium-210

198

The Distribution of Radium-226

200

Anthropogenic Plutonium in the Sea

203

Toward a Model of Metal Transport

205

Distributions of Stable Metals in the Sea

215

Stable Isotope Ratios in Reactive Metals

221

Transport of Iron and Manganese in the Sea

226

Lessons from Controlled Ecosystem Studies

230

Distribution Coefficients

230

Summary230

CHAPTER 5 HOW FAST DOES THE MILL GRIND?
Rates of Vertical Mixing and Sediment Accumulation
Introduction

236

Rate of Vertical Mixing

236

Implication to the Distribution of Radium-226

243

Distribution of Radiocarbon in the Ocean

245

Rate of CDntinental Runoff

247

Sediment Accumulation Rates

253

Radiocarbon Dating

255

Uranium Series Dating

257

Beryllium Dating

264

Potassium-Argon Dating

264

Agreement Among Dating Methods

269

CDmparison of Model and Observed Rates of CaC03 Solution

269

Summary

271

CHAPTER 6 WHAT KEEPS THE SYSTEM IN WHACK?
Control Mechanisms Operating in the Sea
Introduction

275

Phosphate Controls

275

Silicate Controls

280

C:arbon Controls

280

Interactions between the Phosphate and Carbon CDntrols

281

Nitrate Controls

284

Dissolved Oxygen Controls

284

Major Anion Controls

287

Major Cation Controls

289

Possible Causes for Perturbations

300

Recorders of Paleoocean Chemistry301

The Marine Geochemistry of Carbon-13

306

The Uranium Content of Coral

311

Factors Influencing Nutrient Gradients in the Deep Sea

312

Summary314

CBAPTER 7 FREIGHT TRAINS AND FICKIAN CONFUSION
The Movement of Water Through the Deep Sea
Introduction

317

Types of Motion

317

One Dimensional Advection - Diffusion Model

318

Tracers for Diapycnal and Isopycnal Mixing

324

Mixing Rates Based on Radon-222 and Radium-228

325

The Distribution of Helium-3 in the Deep Pacific331

Sources of Deep Water

334

Northern CDmponent Water

335

Conservative Properties of NCW

341

Initial Radiocarbon to Carbon Ratio in NCW

343

Feed for NCW Production

343

Southern CDmponent Water

349

Ventilation of the Deep Atlantic Ocean

359

Ventilation of the Deep Pacific and Indian Oceans

365

The Grand Cycle of Radiocarbon in the Deep Ocean

368

Biological Short-Circuiting372

Temporal Variations in Radiocarbon Production

374

Argon-39

377

Summary379

CBAPTER 8 THE ANTHROPOGENIC INVASION
The Movement of Water Through the Oceanic Thermocline
Introduction

383

Input Functions

383

Hydrology of the Main Thermocline

388

Tritium Distribution within the Thermocline

399

Temporal Trends in Tritium

402

Tritium as a Guide to Deep Water Formation

406

Supplementary Information from Strontium-90

412

Bomb Carbon-14 Distribution within the Thermocline

412

Explanations for Low Equatorial Bomb Carbon-14 Inventories

422

Implications of Equatorial Upwelling to the Tritium Budget

425

An Upwelling Rate Based on the Equatorial C02 Anomaly

427

Helium-3 Distribution in the Main Oceanic Thermocline

432

Purposeful Tracers

438

Summary439

CHAPTER 9 ICE SHEETS AND OCEAN PHOSPHATE
Glacial to Interglacial Changes in Ocean Chemistry
Introduction

444

Temperature and Salinity Changes

446

Formation and Destruction of Organic Materials

451

Changes in CaC03 Storage

456

Evidence for an Early Post-Glacial Lysocline Change

457

Changes in Phosphate Concentration

466

The Combined Evidence from Deep Sea Cores

472

Cause of the Oceanic Phosphate Change479

An Alternate Scenario

480

Wrap Up of the C02 Record

481

The Oxygen Record

484

Glacial to Interglacial Changes in Ocean Mixing Rate

485

Glacial to Interglacial Lysocline Changes

485

Changes in the Distribution of Nutrients in the Deep Sea

491

Summary496

CHAPTER 10 CAN MAN OVERRIDE THE CONTROLS?
The Buildup of Fossil Fuel CO 2 in the Atmosphere and Oceans
Introduction

500

C02 Production in the Past

500

C02 Production in the Future

505

Capacity of the Sea for Fossil Fuel CD2 Uptake

511

Utilizable Capacity - Simplified Calculation

513

Utilizable Capacity - Rigorous Calculation

517

Kinetics of Fossil Fuel CO 2 Uptake by the Sea

524

Numerical Model526

C:rosschecks on the Validi ty of the Numerical Model

534

Prediction of Future C02 Levels

550

Solution of Sea Floor Calcite

552

Summary

564

REFERENCES, CONSTANTS, DEFINITIONS, ABBREVIATIONS AND INDEXIntroduction to the References

569

SUbject Outline for the References

569

Annotated Reference List

574

Frequently Used C.onstants

680

Definitions of Isotope Notations

681

Abbreviations

682

Index

683

Foldout Caption690

etta-pte/!- I
INTERNAL CYCLING AND THROUGHPUTPATHWAYS FROM RIVER MOUTH TO SEA FLOORINTRODUCTION
The sea is a way station for the products of continental erosion.
All substances received by the sea are ultimately passed along to the sediment and rock lining its floor.
The great tectonic forces that continually modify the geography of the earth's
surface eventually push the material buried in this way back above
sea level where it becomes subject to erosion.
Then another trip
through the sea begins.
The greater proportion of the products of erosion enter thesea in particulate form.
They are dropped by the winds to the sea
surface or disgorged by rivers into coastal waters.
These rockfrag'rnents and soil residues are chemically qui te inert.
They
travel only as far as the currents can carry them before reaching
the sea floor.
They play a minor role in the story we have totell.
Of great interest to us are those substances that dissol ve
during erosion and are carried to the sea in ionic form.
They
constitute the sea's salt.
As long as they remain dissolved,
gravi ty cannot influence them.
Al though the processes at work inthe sea ul tima tely " reprecipi tate" these ions, most cons ti tuentsof sea sal t wander through the sea long and far before becoming
entombed in the sediment.
The temporary entrapments in particulate matter experienced by these wandering ions will prove of particular interest to us.
The composition of sea salt reflects not only the relative
abundance of the dissolved substances in river water but also theease wi th which a given substance becomes entrapped in the sediments.
Sodium, for example, is both abundant in the dissolved
matter in ~ivers and sparingly reactive in the sea. This combination is reflected by its high concentration in sea salt. Calcium,
al though even more abundant in river water than sodium, is an important ingredient in the shells of marine organisms.
Because ofthis special mode of entrapment, the abundance of calcium in sea
water is far lower than that of sodium.
Many components of sea salt show little variation in concentration within the sea. Others show very large changes in concentration from place to place.
As we shall see, these differences
are largely the result of cycling by organisms.
Plants live onlyin surface waters, from which they extract certain constituents1

needed to construct their tissues.
While much of this plant matter is consumed by animals living in the surface ocean, some insolubles and indigestibles (i.e., fecal matter) move into the deep
sea under the influence of gravi ty - and so the life cycle leads
to chemical segrega tiona
Des true tion of plant rna tter occurs, on
the average, at greater depths than formation.
The interaction ofthis life cycle with the large-scale water circulation pattern inthe sea resul ts in inhomogeneities in the distribution of these
species, not only within the sea itself but also in the sediments
deposited on the ocean floor.
One of the aims of this book is to point out the factors that
influence the average concentrations of the various components ofsea sal t and the fac tors tha t produce chemical inhomogenei ties
wi thin the sea and its sediments.
The approach might be termed"inverse chemical engineering".
The ocean is a great chemicalplant that processes the dissolved matter added from rivers anddispenses it as sediment.
Unlike most chemical plants, the sea
has no advance operational blueprint.
As chemical oceanographers,
we wander through the plant measuring inputs, outputs, and internal compositions - trying to reconstruct the missing design.
As inmost chemical plants, the two critical features of the ocean are
the manner in which the ingredients are mixed and the manner inwhich the ingredients react wi th one another.
Oceanic mixing isaccomplished by a complex system of currents and a host of turbulent eddies.
Many of the chemical reactions are catalyzed by the
enzymes in living organisms.
Thus any study of the chemistry ofsea water is heavily dependent on knowledge derived from physicaloceanographic and marine biologic studies.
In this book, we will focus on the first-order processes
operating within the sea.
Until these are mastered, it is fruitless to proceed to the more complex "details". With this in mind,
let us turn our attention to our first subject - the grand chemical balance existing in the sea.

DEPTH PROFILES OF SEA SALT COMPOSITION
The salt dissolved in sea water is remarkably uniform in its
major constituents.
This fact has greatly simplified the task ofthe physical oceanographer interested in mapping water density
patterns within the sea.
He needs only to measure the water temperature, the water pressure and one major property of sea sal t(for example, the chloride ion content or the electrical conductivity) to make a very accurate estimate of the in situ density ofa given sample of sea water.
This task would be-extremely complex
if the composition of sea salt were more varied.
Yet it is fortunate, too, that the compositional constancy ofthe major components of sea salt does not extend to all the minor
components.
If this were not the case, the oceanographer would
lose one of his most powerful tools, for the varia tions in theminor consti tuents of sea wa ter bear important clues regarding
mixing, biological, and sedimentary processes taking place within
the sea.
They are the tracers about which this book is written.
The major ion matrix of sea salt consists of the following
constituents:
Chlorine in the form of CI- ion; sulfur in the form2

of SO 4 = ion· and magnesiwn, potassium, calcium and sodium in theform of Mg+q., K+, Ca++, and Na+ ions, respectively.
These sixions dominate sea salt; their ratios, one to another, are very
nearly constant.
In fac t, only calciwn has been shown to varymeasurably from place to place in its ratio to the other five cons ti tuents.
Al though this cons tancy also extends to many of thelesser components of sea water (boron, bromine, fluorine, uranium,
cesium, and others), it does not extend to all of them.
The best known of the varia tions in the composi tion of sea
salt arise from the removal of constituents from surface sea waterby plants and the subsequent destruc tion of plant-produced particles after downward movement.
Deep water masses are richer in theconstituents utilized by plants than surface water is. If the ocean were sterile, the chemical composition of sea sal t would bealmost perfec tly uniform.
The main differences in composi tion
would be those resul ting from the transfer of gases between theatmosphere and surface waters of differing temperature would
exis t.
Plants live only in surface water, where there is enoughlight to permi t photosynthesis.
By the time the components of
their debris are returned to the ionic form as a resul t of tissue
oxida tion or hard part solu tion, downward movement under the influence of gravi ty or migra ting animals has occurred.
It is notsurprising, then, tha t the primary chemical differences observedin the ocean are all of the type just mentioned: deep water is enriched relative to surface water.
The only major exception is
dissolved oxygen.
Figures 1-1 and 1-2 show the dis tribu tion wi th depth of potential temperature*, of salinity and of the concentrations of six
biologically utilized sea salt constituents at a station in the
North Pacific Ocean.
In all cases, the most dramatic change occurs in the upper thousand meters of the wa ter column.
This socalled main thermocline constitutes the zone of transition between
the warm surface wa ters and the cold deep wa ters.
The nu trient
species nitrate (N03-), silicate (H 4 Si0 4 ), and total dissolved inorganic carbon (L:C02 = C02 + HC03- + C03=) show the deep water
enrichment mentioned above.
Alkalinity and barium show patterns
similar (but not identical) to that of silicate. Dissolved oxygengas, by contrast, shows a depletion.
Unlike N03- and L:C02, which
are released during respira tion, 02 is consumed during respiration.
Its pattern is also complicated by the fact that cold waters descending from the surface carry with them more dissolved
gas than descending warm waters.
The salinity minimum at about
600 meters depth is genera ted by the la teral intrusion of intermedia te wa ters.
Keep in mind tha t these frac tional salini ty differences are quite small compared to the fractional differences in*Sea water is heated by compression (i.e., about O.loC/km) as
it descends from the surface.
Potential temperatures are corrected for this heating.
Hence the potential temperature is
the temperature a water sample would have were it returned tothe sea surface.

3

SALINITY (%0)

POTENTIAL TEMPERATURE (Oe)

00

10

20

54r·~O:::::34~.~2:::!::3::41·:4~=3~4~.6~~

30

E 2

E

.:a::

.:a::

~
~

J:

UJ

lLJ

~

a..

a..
o

o 4

6 ......-

61-.....- ......_ . . & . . - 1 0 -. . . .-......- - - -......- - ' - -.....--'.....- -300--~---~

E

~

-

l-

I-

~

~

a..
w
o

e.

w
Q

Figure 1-1. Plots of temperature, salinity, dissolved oxygen content and ni trate content as a function of water depth at GEOSECS
station 214 in the North Pacific (32°N, 176°W).
Potential temperature rather than the temperature measured in situ is given.
The salini ty profile in this region of the ocean--sh"ows a pronounced minimum at a depth just over 600 meters.
This intermedia te water .C as it is called) forms in the northern Pacific andsinks and flows laterally beneath the waters of the warm temperate
ocean.
The level of this minimum is shown in the other diagramsby a dashed line.
Dissolved oxygen is utilized by animals andbacteria Ii ving in the deep sea to "burn" the organic debris falling from the surface.
Thus all deep waters are deficient in oxygen compared to the amount they recei ved from the atmosphere before descent. As oxygen is consumed, ni trate is produced.
Thus
the shapes of the profiles of these two properties have an inverse
relationship. These resul ts were obtained as part of the GEOSECS
program (425).
The loca tion of this s ta tion can be seen on thefoldout map.

4

ALKALI N ITY (~)

TOTAL I NORGAN I C CARBON (,u.mol/I<<;ll

o

2000

2200

2300
2400
2500
Or---......--....--......
--..---..~~

2400

-E 2
~

6"--.....-

6"'--......- - -.....---'--.......-----'
SILICATE (pomallkg)

50

E 2

~

J:
l-

a.
l.IJ
o 4

100

150

.....- - - -.......- .....................
BARIUM (nmal/kg)

100
150
Or-'lf'-.,...-...,...--,.-......
-......,r-----.

E

2

~

J:

I-

a.

IJJ

0

4

---"-----.-.10-.............--

6----.....

Figure 1-2.
Plots of total dissolved inorganic carbon, alkalinity, dissolved silicate, and dissolved bariwn as a function ofwater depth at GEOSEGS station 214 in the North Pacific (32°N,
176°W).
As in figure 1-1 the dashed line represents the depth of
the salini ty minimum.
The alkalini ty of the sea changes mainly
because Ca is ext rae ted from surface wa ters by organisms to form
GaCO 3 skeletal material.
Much of this CaGO 3 dissolves after .falling to the deep sea.
The depth profiles of these cons ti tuentsdiffer from one another and from that of ni tra te because they arecarried down in different forms; C and N in organic debris, C and
Ca in CaC03, Si in opaline silica, and Ba in some as yet unidentified form.
The depth spectrum of the destruction of these various
phases is not the same. The skeletal matter reaches, on the average, greater depths before dissolVing than the organic debris
reaches before consumption.
This gives rise to deeper maxima foralkalinity, silicate and barium than for nitrate and carbon.
These results were obtained as part of the GEOSECS program (425).

5the concentrations of the nutrient species.*

CONSTITUENT CLASSIFICATION
Roughly one quarter of the 92 elements will be considered inthis book. They are shown in table 1-1, grouped as they appear in
the chemist f s periodic table.
The first column contains a groupof elements known as the alkali metals; they have a valence of +1(that is, in sea water each atom loses one electron and becomes an
ion with a single posi ti ve charge).
The alkaline earths occupy
column two. These elements have a common valence of +2 (two electrons are given up upon solution in sea water).
The noble gases
are in column eight.
They do not undergo chemical reactions in
sea water, but remain in gaseous form.
Column seven contains elements with the common property of accepting an extra electron (oneof the electrons released by the alkali metal and the alkaline
earth metal atoms).
These elements are present in sea water witha charge of -1.
The elements in columns three, four, five and six
all combine with oxygen, and sometimes with hydrogen, to form multiatom ions.
In most cases, they form negatively charged ions; ina few cases, they form neutral groupings.
Those dissol ved uni ts
with negative charges are known as anions; those with positive
charges, as cations.
In addition to the elements listed in table
1-1 the marine chemistry of a number of the transition metals(iron, manganese, copper, zinc, nickel, and cadmium) and heavy
metals (uranium, thorium, protactinium, polonium, lead and plutonium) will be considered.
The chemical constituents of sea salt can be divided into
three major categories:
the biolimiting constituents (those which
are almost totally depleted in surface water); the biointermediate
constituents (those which are partially depleted in surface water); and the biounlimited constituents (those which show no measurable depletion in surface water).
The six known biolimiting constituents are nitrate (N03-),
phosphate (HP0 4 =), silicate (H 4Si0 4 ), zinc (Zn++), cadmium (Cd++)and germanium (H4Ge04).
Plant activity (and, in the case ofsilicate, animal activity as well) is efficient enough to extract
these six constituents almost completely from surface water.
The
deep sea constitutes a large reservoir for these constituents.
When deep water is returned to the surface, these constituents
become available to photosynthetic organisms. They are fixed into
particulat.e material and are then carried by graVity back to thedeep sea, where they are largely returned to solution.
The question naturally arises as to how more than one chemical ingredient can limit plant growth.
The answer li es in thefact that many types of plants live in the ocean.
Different types
have different chemical needs and abilities. For example, a groupof plants called diatoms need H4 SiO 4 to manufacture the opaline*A 0.6 per mil salinity change is equivalent to a two percent
change in salt concentration.
A per mil (%0) is a part per
thousand; just as a percent (0/ 0 ) is a part per hundred.

6

Table 1-1.
Abbreviated periodic table showing the elements whose marine chemistry are
discussed in this book.
For each element the important ionic and molecular forms foundin sea water are noted.
For those elements whose distribution within the sea is sufficiently well understood, the designations biolimiting, biointermediate, and biounlimited are given.
~ot shown are the transition metals (Fe, Mn, Cd, Zn, Ni, and Cu) andheavy metals (U, Th, Pa, Po, Pb, and Pu) which are discussed in Chapter 4.

(8)

(1)

""'-J

Hydrogen
H2 O
H+
( 4)
(2)
(3)
Unlimi ted
CarbonLithium Beryllium Boron
Be++
Li+
HC03H3 BO 3
H4 B0 4 C03=
Unlimited CO2
Intermediate
SodiumMagnesium Aluminum Silicon
Mg++
Na+
Al(OH)4- H4 Si 04
Limiting
Unlimited Unlimited
Al(OH)3°
Potassium Calcium
K+
Ca++
Unlimited Intermed.
Rubidium Strontium
Sr++
Rb+
Unlimited Intermed.
CesiumBariumBa++
Cs+
Unlimited Intermed.
RadiumRa++
Intermed.

Selenium
Se03=' Se04=
Intermediate

HeliumHe( 5)

Nitrogen
N2
Unlimited
N03Limi ting
Phosphorus
HP04=
H2 P0 4Limiting(6)
OxygenH2 O
Unlimited
02
Intermed.
Sulfur
S04=
Unlimited(7)
Unlimited
Fluorine Neon
FNe
Unlimited Unlimited
Chlorine Argon
ClAr
Unlimited Unlimited
BromineKrypton
BrKr
Unlimited Unlimited
Xenon
Xe
Unlimited
Radon
Rn
Unlimitedcapsules within which they dwell.
Thus, the dissolved silica in
sea water limits the extent to which diatoms grow.* Another plant
type, the blue-green algae, has the capability to produce forms of"fixed lf nitrogen (i.e., N03-) from dissolved nitrogen gas (i.e.,
N2).
Thus, the growth of these plants need not be limited by theavailabili ty of NO 3-.
All plants need the element phosphorus.
Hence, phosphorus is the ultimate limiting nutrient for plant
growth.
The biological roles of Zn and Cd are unknown.
Plantsmay inadvertently incorporate them in their quest for phosphorus.
A consti tuent is biounlimi ted if its ratio to total salt inboth surface and deep sea water samples are equal within measurement error. The elements currently classified as biounlimited are
sodium, potassium, rubidium, cesium, magnesium, boron, sulfur,
fluorine, chlorine, bromine and uranium.
If more accurate analyses could be made, one or more of these elements might be eliminated from the biounlimited category.
We are certain of eight biointermediate elements:
calciumeCa), strontium (Sr), nickel (Ni), copper (Cu), selenium (Se),carbon (C), barium (Ba), radium (Ra).
The Ca content of surface
water salt is about one-half percent less than that of salt indeep water.
The C content of surface water salt is 15 percent
less than that of deep water, and the Ba content of surface water
salt is 70 percent less than that of deep water.
In order to fit into the classification adopted here a constituent must be recycled within the sea several times before removal to the sediments.
Some elements are so reacti ve that they
pass through surface water only once.
They become firmly fixed
onto particles which fall through the sea to the sediment.
These
constituents will be referred to as noncycling.
Iron, thorium,
protactinium and lead fall in this category.

THE CHEMICAL COMPOSITION OF MARINE ORGANIC MATTER
Three major types of particles fall from surface water into
deep water:organic tissue, calcium carbonate (CaC03), and opaline silica (Si02). All plants and animals produce organic tissue.
Plants extract the ingredients for this tissue from the dissolved
salt in surface sea water.
Animals reuse these ingredients by
devouring plants and other animals.
Roughly, 90% of the organic
matter eaten by animals is consumed in the production of energy.
*At t'he time we were completing this book a paper by Froelichand Andreae of Florida State University was published showing
an almost perfect covariance between the concentrations of theelements germanium and silicon for samples from various depths
at a North Pacific GEOSECS station.
While it is tempting tocone lude from this that Ge and Si have identical marine chemistries (i .e., they act as a single element), Froelich has
further unpublished data suggesting that the situation is notnearly so simple (378).

8

The products of this "combustion" are returned to the sea in ionicform.
Thus only about 10% of the organic matter survives eachstep in the food chain.
Of this some is rejected in fecal pellets
which fall toward the sea floor.
Many microscopic plants andanimals also produce hard parts consisting of CaC03 or 3102. Hardparts made of CaCO 3 are produced by cocco11 thophor1da (plants),
foraminifera (animals), and pteropods (animals).
Hard parts madeof 8i02 are produced by diatoms (plants) and radiolarians (animals).
Photographs of the hard parts produced by these organisms
appear in figure 1-3.
The chemical composition of the organic soft tissue formed byplants is thought to be relatively constant:for every atom of P
in this tissue there are on the average about 15 atoms of N and on
the average 105 atoms of C.
The approximate ratios of these sameelements dissolved in deep sea water are 15 atoms of Nand 1000
atoms of C for every atom of P.
When a batch of typical deep
Pacific water returns to the surface, plants extract phosphate andnitrate until both of these constituents have been depleted.
Neglecting for the moment the carbon lost to CaC03, only 105 of
the 1000 atoms of C w111 be used.
Roughly 90 percent of the carbon will remain in dissol ved form.
In the absence of addi tional
phosphate and nitrate, plants have no use for this remaining dissolved carbon.
Animal tissue has roughly the same P/N/C ratio as
plants.
This makes one aspect of ocean chemistry qui te simple.
To the first approximation, parcels of sea water differ in theirP, N, and C contents to the extent that living organisms haveremoved or added these elements in a reasonably fixed ratio.*
One of the mysteries of sea water chemistry concerns the PIN
ratio.
When deep water upwells to the surface, by the time all of
its dissolved phosphate has been consumed, by chance or by design,
so has its dissolved nitrate.
Why Nand P are present in sea
water in the same ratio as in organisms is a little like thechicken and the egg problem.
Do organisms use Nand P in the 15/1
ratio because this is the ratio in which these elements occur in
sea water, or do organisms maintain this 15/1 ratio in the sea?
The answer is almost certainly that organisms drive the ratio of
NIP in sea water toward the 15/1 biochemical ratio. The mechanism
must be ecologic.
If the amount of nitrate in the sea were todrop below the amount required to maintain this ratio, then species which are capable of fixing molecular ni trogen (i. e. con-*This ratio is known as the Redfield ratio in honor of Alfred C.
Redfield who demonstrated the relationship between organism composition and ocean chemistry. Although the C:N:P proportions of 105:
15:1 are widely used the exact values are still quite uncertain.
For example, analyses of material from surface water yields a range
in C:P ratios of 50 to 200 (692).
A similar range is observed in
fresh waters.
Thus, while representing a fair average compositionof the micro-organisms in the oceans, sizable differences are to be
expected from place to place in the sea.

9

I 3fL I

(0 )

(b)

(c)

(d)

Figure 1-3. Photographs of hard parts produced by (a) coccolithophorida" (b) foraminifera, (c) diatoms" and (d) radiolarians.

10verting N2 to N0 3 ) would gain an advantage and would begin to produce nitrate faster than it was being destroyed within the sea.
On the other hand, if an excess of N03- relative to the amountneeded to match the biochemical ratio existed within the sea thenthose species that do not fix nitrogen might be expected to havean advantage.
Presumably the ecology of the sea adjusts until abalance is achieved such that the rate of nitrification justmatches the rate of denitrification.
This balance is struck when
the NIP ratio in sea water is quite close to that in the microorganisms which dominate the life cycle in the sea.
The plot in
figure 1-4 of N0 3 versus P0 4 for waters from throughout the Atlantic Ocean demonstrates the remarkable constancy in the ratio ofthese two constituents in the sea.
Carbon is also used by organisms to construct their calcium
carbonate hard parts.
In CaC03 there is one C atom for every Ca
atom.
In sea water there are 20 C atoms for every 100 Ca atoms.
Thus, the generation of CaC0 3 depletes carbon about five times
more rapidly than it does calcium.
On the average, for every four C atoms that fall from thesurface sea in the form of organic tissue, one C atom falls in theform of CaC0 3 .
Thus for every P atom in the composi te of theorganic and CaC03 debris that fall toward the deep sea there are
131 C atoms, 105 of which are part of organic material and 26 arepart of CaC03. Accompanying them, of course, would be 26 Ca atoms
paired off with the carbon atoms in the CaC0 3 (see table 1-2).
In
deep sea water, on the average, for every atom of P there are 1000
atoms of C and 5000 atoms of Ca.
Upon total depletion of Nand P
only 131 (13 percent) of the 1000 C atoms originally present indeep sea water are used to form organic tissue and CaCO 3.
Forevery 5000 Ca atoms available, organisms consume only 26 (i.e.,
only 0.5 percent of the available Ca).
This explains why calcium
is almost but not qui te constant in the sea salt of the world
ocean.
For every atom of P in the deep ocean there are about 50
atoms of Si.
Since surface-dwelling organisms use essentially all
the available silica, we can add the appropriate amount of opal tothe debris.
Because the amount of opal containing 50 Si atoms has
nearly the same weight as the amount of CaC03 containing 26 C
atoms and as the amount of organic tissue containing 105 C atoms,
the average batch of falling particulate matter consists of equal
weights of its three major constituents.
At any particular pointin the ocean, their proportion may vary significantly from thisaverage because of local ecologic differences in the overlying
surface water and because the hard parts (CaC0 3 and 8i02) fall togreater depths before destruction than does the organic tissue.
As organic tissue sustains animal life it is a rapidly consumed
commodity; hard parts, on the other hand, have no energy value andare left to undergo gradual inorganic solution.
Indirect evidence leads us to believe that barium is removed
from sea water largely as barium sulfate (BaSO 4) •In order toexplain the observed depletion in surface water, only 1 atom of Ba
need be removed for every 3000 atoms of C; thus the presence of
BaS04 in the debris can easily go unnoticed.

11

Figure 1-4. Nitrate versus phosphate scatter diagram for watersfrom throughout the Atlantic Ocean as determined during the GEOSECS program (685).

3r---.---.-----.--...--.,.---...----.r--........

~

o

0..

10

20

40

30

N0 3 (,umol/kg)

Table 1-2.
"Redfield" ratios in particulate debris falling fromthe surface toward the deep sea compared with the elemental composition of average surface and deep water.
These compositions
are termed Redfield ratios in honor of Alfred Redfield who did theearly work on this subject (346).
For comparison actual P:N:C
ratios measured in material sieved from surface waters are shown.
These latter results were taken from a summary by Goldman (692).
P: N:

C :

Ca:Si

0: 0
26:50
26:50

Particulate Debris.

Soft Tissue
Hard Parts
Composite

1:15: 105:
0: 0: 26:
1:15: 131 :Sea Water

Deep
Warm Surface

1:15:1000:5000:50
0: 0: 869:4974: 0Western North Atlantic (Jan. '62
Western North Atlantic (Apr. '62)
Equatorial Atlantic
North Pacific Gyre
Central North Pacific12

P: N: C
1: 5: 59
1:13: 68
1:21:163
1:29:410
1:18:152

The element radium has chemical affinities almos t identicalto those of barium, and the element strontium (Sr) has affinities
quite similar to those of calcium.
Thus the formation of BaSalt
and C:aCD 3 depletes radium and strontium as well as barium andcalcium in surface water.
However, because the Sr/Ca ratio in
most marine CaCa 3 is only about one-fifth that in seawater, theexpected depletion of Sr (about 0.1%) lies within the limits ofmeasurement error.
Strontium has been shown, however, to be
deficient by about 1% in surface water.
The likely explanation
for this larger than expected deficiency is that the mineralcelestite (SrSa Lt ) is produced by some organism living in surface
waters.
As sea water is highly undersaturated with respect tothis mineral phase, it redissolves after only a few hundred metersof descent.

COMPOSITION OF PARTICULATE MATTER CAUGHT IN SEDIMENT TRAPS
The most direct way to study the composi tion of marine particulate debris is to deploy and recover traps designed to catch
this material as it settles through the water column. This technique has been used effectively in the open ocean only during thelast several years.
Even now only a few of the resul ts of suchstudies have appeared in the li terature.
The late bloom of thisline of research parallels the development of large deep sea
moorings by physical oceanographers.
The big international programs ini tiated in the 1970' s led to the design and testing ofbottom moored open ocean devices on which it was possible tomount current meters arrays covering the entire water column.
Such moorings can be called back after many months of service bymeans of sonic systems.
The weight which carries the instrument
string to the bottom and holds it in place there is released by asound message sent from the recovery ship; the buoyant spheres
attached to the arrays then float the whole system back to thesurface where it can be hoisted onto the recovery ship.
Honjo, at Woods Hole, first mated this technology with mode~nsediment traps.
His traps were large (cross section about 1 m )as he wanted to trap enough material to permit the myriad of chemical, radiochemical, stable isotope, paleontologic and sedimentologic analyses demanded by such studies.
To date Honjo has deployed his traps in the north temperate Atlantic, the north temperate Pacific, the equatorial Atlantic and in the equatorial Pacific close to Panama.
His efforts have paid off handsomely!
Inthis chapter we will mention some of the chemical and sedimentologic resul ts.
In chapter 4 we will discuss the radiochemical results obtained on these samples.
The resul ts for the sediment traps deployed by Honjo in the
open ocean are summarized in table 1-3.
Wi th increasing depth
continental detritus constitutes an ever greater fraction of thematerial caught.
While this change is in part the result of destruction of the biogenic components as they fall through thewater column, it is also in part the result of increasing flux of
detrital material with depth in the sea. This increase is the result of resuspension of detritus from the "bottom" and "sides" ofthe ocean.
Silicate detritus dominates this resuspended material

13

Table 1-3. Sediment trap data from the Atlantic and Pacific Oceans.
The depths at which the traps were deployed are given at the top ofeach column. These results were obtained by Honjo and his coworkers
at the Woods Hole Oceanographic Institution (541).

ATLANTIC

2°S-_30 0 N_
1.0 lan 4.0 krn

0.4 kIn

1.0 kIn

3.8 lan

5.1 kIn

Clay (%)
CaCO 3 (%)
Opal (%)
Organic Matter (%)

7
63
10
20

16
55
12
17

25
57
9
10

31
49
9
11

7
59
7
27

12
68
9
11

Org C (moles/m 2yr)
CaC03 (moles/m 2yr )Opal (moles/m 2yr )

0.21
0.15
0.035

0.12
0.09
0.031

0.05
0.09
0.023

0.05
0.08
0.023

0.074
0.040
0.008

0.027
0.043
0.010

Clay (gm/cm 2 10 3yr)
CaCO 3 (gm/cm 2 10 3yr )Opal (gm/cm 2 10 3yr)

0.2
1.5
0.2

0.3
0.9
0.2

0.4
0.9
0.1

0.5
0.8
0.1

0.005
0.40
0.05

0.08
0.43
0.06

C:N (mole: mole)
N:P (mole:mole)
C:P (mole :mole)

7.6
23.7
180

8.5
31.8
270

10.9
16.5
180

8.4
23.7
199

5.1
17.7
90

9.3
11.9

PACIFIC

III

15°N

Clay (%)
Opal (%)
CaCO 3 (%)
Organic Matter (%)

0.4 lan2
3
35
60

100 krn
3972
16

2.8 kIn215
68
14

4.3 kIn3
15
72
11506 krn
4
21
61
14

Org C (moles/m 2 yr )CaC03 (moles/m 2 yr)
Opal (moles/m 2 yr)

0.108
0.014
0.003

0.017
0.019
0.016

0.033
0.041
0.016

0.027
0.042
0.015

0.020
0.024
0.014

Clay (gm/cm 2 10 3yr)
CaC03 (gm/cm 2 10 3yr)
Opal (gm/cm 2 10 3yr)

0.01
0.14
0.01

0.01
0.19
0.02

0.01
0.41
0.09

0.02
0.42
0.09

0.02
0.24
0.08

C:N (mole: mole)
N:P (mole:mole)
C:P (mole:mole)

8.4

9.6

-

10.2
26.1
266

10.0

-

10.6
29.1
309

14

-

because the biologic components which reach the deep sea floor arelargely destroyed by respiration and solution.
Also shown in this data set is evidence that organic tissue
suffers greater destruction as it falls through the water column
than do CaC0 3 and opal.
Honjo has shown that the flux of calcite
is amost constant with depth.
The changes in the CaC03 flux thatare seen are result of aragonite solution.
The C/P and NIP ratios in the sediment trap material are sUbstantially higher than the Redfield ratios discussed above.
Whether this is an artifact of microbial activity in the traps
or analytical procedures or a true measure of the time flux ratios
has yet to be demonstrated.
A SIMPLE MODEL FOR BIOLOGICALLY UTILIZED CONSTITUENTS

Now let us be a bit more quantitative about the operation of
the ocean's internal cycles.
Geochemists who study the sea findit convenient to treat this immense system as a series of well
mixed reservoirs. For our rather elementary look at the ocean, we
will divide it into just two such reservoirs - the warm waters andthe cold waters.
We do so because the main obstruction to mixing
within the sea is the density difference between the thin skin ofwarm surface water which covers the equatorial and temperateregions of the ocean and the cold wa ter found at the surface inthe polar regions and throughout the deep sea.
The zone separating these two maj or wa ter types (referred to hereafter as theoceanic thermocline) lies between 70 meters (the average thicknessof the wind-stirred surface mixed layer) and 1000 meters (the beginning of the deep wa ter regime).
As the poles are approached,
the oceanic thermocline rises to the surface and provides a horizontal separation between polar and temperate surface waters.
Thebulk of the ocean's plant life lives in the upper, sunlit portionof the warm water mass.
Al though the cold reservoir is also
lighted by the sun at its polar "ou tcrops", the area of this exposure is relatively small.
This simplification of the ocean allows us to quantify its basic processes ra ther easily.
Once weunderstand the two-box model shown in figure 1-5, we will have afirst-order view of how the ocean operates.
We will make further simplifications.
First, we will assume
that the only way a cons ti tuent is added to the sea is by runoff
from the continents.
We will neglec t other means of entry, suchas hydrothermal
activity on the ocean floor and ground waters
seeping out along the continental margins.
Second, we will assume
that the only way biologically u til ized cons ti tuents are removed
from the ocean is by the fall of organism-produced particles tothe sea floor.
Since such particles are not entirely destroyed by
predators, scavengers, bacteria, and corrosive waters, a small
frac tion of this debris is permanently buried in the sediment.
Thus rivers add rna terial to the sea and the particles formed insurface waters fall to the bottom, removing material from the sea.
We know that the biologically used constituents are being replacedin the ocean wi th great regulari ty (at least geologically speaking).
The sea gains from rivers (and loses to its sediments) its
total content of most of the biologically important cons ti tuents

15

,

t:·

....
.:::

Input fromRivers0·.···.

(Jt

IIrjv~:r x 0

Net Evaporation
::::
:::_1

SURFACE OCEAN

}I

C'~'

UPw-~~fllng

Down'~:elling

Water

W~:~er

.~~~

=::

en

..

:::

.....

......

SURFACE OCEAN

:~::I

.'.

:t

·t

},

~)
:::

IImlxCsurf!e

m~i

":;:

7i~ixCdeep

~~

:::

~i~

Preserv'ei{ In Sed i ment

'Y

Figure 1-5. The diagram at the left is the two-box model showing the major fluxes ofwater and of biologically produced particles in the sea.
Straight arrows indicate
the fluxes of water; wavy arrows indicate the fluxes of particles.
The diagram at
the right shows the same model with the fluxes labeled in accord with the parameters
defined in the text.

in time periods between 10,000 and 1,000,000 years.
Recent studies or the chemistry of the waters that are circulated between the deep sea and the hot basaltic rock which is continually being added to the expanding ridge c res ts suggest tha t
the basalt-water interactions act as a sink for some constituentsof sea salt (i. e., they are lost from the sea water as it percolates through the hot rock) and as a source for others (i.e., thehea ted wa ter leaches rna terial from the basal t and carries it to
the sea).
As these interactions do not appear to have an important impact on the budgets of the biologically important constituents we neglect the effects of this process in our first ordermodel.
In our model, we will further assume that the operation of
the ocean is at steady state.
By this, we mean that the rates of
inpu t and au tpu t of any cons ti tuent from the ocean as a whole aswell as from the warm or the cold reservoirs have remained constant for a sufficiently long time that the concentrations at any
point in the sea are not changing with time.
The system has stabilized so that gain just balances loss.
The content of any cons ti tuent is being steadily renewed, but its amount is not changing.
The si tua tion is much like the flow of bot tIes through a
Coca Cola bottling machine.
Empty bottles are steadily added at
one end of the conveyor system; full bottles emerge at the other.
The number of bottles on the conveyor is always the same.
Photographs taken at various times would look alike.
As biologically active consti tuents pass through the ocean,they participate in its internal cycles.
Warm surface waters receive their supply of dissol ved cons ti tuents from two sources:
river water and water from the deep ocean which exchangs with surface water.
If the concentration of a given constituent in thesurface ocean is to remain constant, then loss with downwelling
water and with settling particles must match these two inputs.
If, as is the case for biologically active constituents, theconcentration of a constituent in warm water is less than it is incold wa ter, then downwelling alone is not adequate.
Since thewa ter exported to the cold reservoir carries away less of eachconstituent than the upwelling water brings back, particulate loss
must account for the difference.
We can rather simply estimate the amount of any element thatdrifts down as a constituent of particles.
The fluxes to be considered in such a calculation are shown in figure 1-5. The volume
of river water entering the ocean each year is vriver; the volume of deep wa ter ris ing to the surface each year is 'Vmix; theconcentration in moles per liter* of the element in average deep
wa ter is Cdeep; the concentra tion of the element in average surface wa ter is Csurface; and the concentra tion of the element inaverage river wa ter is Criver.
The amount of an element entering the surface ocean from river water must be vriver Criver(the volume of water per year mul tiplied by its concentration of
the constituent of interest) plus the upwelling contribution
23*There are 6 x 10
atoms in a mole.
One mole of any element has aweight in grams equal to the atomic weight of the element.
Thus amole of C weighs 12 g; a mole of Si, 28g; and a mole of H, 1 g.

17

VmixCdeep (the volume of upwelled water mul tiplied by its concentration of the constituent of interest).
This input to the surface ocean must be balanced by the loss
fluxes.
The amount of an element leaving the surface via downwelling is vrnixCsurface.
For every unit of water entering theocean from rivers, one uni t is los t by evaporation followed bycontinental rainfall.
As this evaporated water does not carrysignificant amounts of the constituents of concern to us in thischapter, it does not participate in the removal process.
Theparticle flux, P, can be calculated by summing the two inputs tothe surface reservoir and sUbtracting from this total the amountbeing carried to the deep reservoir by downwelling water:
P = vriverCriver + vrnixCdeep - vmixCsurface

1-1

If we define "g" as the fraction of a given constituent
reaching the surface reservoir removed in particulate form, thenthe balance between loss and gain for the surface reservoir can bereexpressed as follows:g(VriverCriver + VrnixCdeep)
vriverCriver + VrnixCdeep - vmixCsurface

1-2

SolVing for g,
vriverCriver + vrnixCdeep - vmixCsurface
g

1-3

or

g

vmix

Csurface

VriverCriver

1 -

1-4

Cdeep

1 +

v rive r Crive r

As we shall see in chapter 5, in the present-day ocean the ratioof vrnix to vriver is about 30; that is, 30 times more water isadded to the warm reservoir by upwelling than by river runoff.
Thus:

30
g

1

-

Csurface
Criver

1 + 30

1-5

Cdeep
Criver

For the constituent phosphorus, Cdeep/Criver is about 3 andCsurface/Criver is 0.15 (on the average, rlver water carries
three times less phosphate per liter than deep sea water, and seven times more than surface sea water). The corresponding value ofg is 0.95 (95 percent of the phosphate reaching the surface ocean18

is carried away by falling particles).
On the average, the amount of any given constituent leaving
the sea, L, must be matched by that being added by rivers:
L = vriverCriver

1-6

Thus, if the removal is accomplished by the burial in the sediments of particles which survi ve des truction, then the fraction,f, of a given constituent carried to the deep sea in particles
that survive destruction,must be equal to the river input:
or

1-7

fP = vriverCriverf(VriverCriver+VmixCdeep-VmixCsurface)=VriverCriver

1-8

or
1

f

1-9
Cdeep

1 +

vrivercsurfaceJ[ Criver

CriverAnd since vmix/vriver = 30:
f1=

1 + 30Cdeep[ Criver1-10
Csurface]
Criver

For phosphate, f turns out to be about 0.01.
Only 1 percent of
the phosphorus carried to the deep sea by falling particles is notrecycled.
Now if we mul tiply the parameter f by the parameter g, we
obtain a rather interesting piece of information.
Since g tells
us the fraction of a cons ti tuent that reaches the surface and isremoved in particulate form, and f tells us the fraction of thoseparticles that survive destruction, the product f x g gives us thefraction of an element that is removed per oceanic mixing cycle(that is, per transfer from the cold to the warm reservoir).
f

1

x g'=

1 +
f

x g

vmix

Cdeep

vriver Criver1
Cdeep
1 + 30
Criver

1-11

For phosphate, the product of f x g would be 0.01 (for f) x
Thus 1 percent of the phosphate in

0.95 (for g), or abou to. 01.

19

the ocean is lost to the sediment during each mixing cycle.
If wewait the length of time necessary for upwelling to recycle all the
wa ter in the ocean through the surface ocean layer and allow theplants to extract the phosphorus, a loss to the sediment of 1 percent of the phosphate in the ocean will ensue.
As we will see from the 14 C dis tribu tion in the ocean (to be
discussed in chapter 5), the time required for one mixing cycle isabout 1000 years.
Every 1000 years, the entire amount of any
given constituent in the ocean is sent through the surface water
mill. To find the time, T, required to remove an amount of a cons ti tuent equal to tha t stored in the sea, we divide the mixing
time, Tmix, by the product of f x g.
In equation form:
T

=

Tmix
fg

1000
fg

years

1-12

Since phosphate is removed at the rate of 1 percent every
mixing cycle (f x g = 0.01), the average lifetime of a P atom inthe ocean mus t be abou t 100, 000 years.
The resul ts of these calculations for the phosphorus are summarized in table 1-4.
Inwords, a typical P atom, upon release to solution from some sedimentary or metamorphic rock by erosion, is carried by rivers to
the sea.
It then goes through an average of 100 oceanic mixing
cycles:
100 times surface dwelling organisms fix it into a particle tha t sinks and is des troyed in the deep sea.
Each time , itwaits in the dark abyss about 1000 years before being sent back tothe surface.
On the average, during the hundredth cycle, the particle bearing the P a tom survives des true tion and is trapped in
the sediment.
So our P atom makes 100 round trips of 1000 years
each during its stay in the oc ean.
It then becomes part of thesediment, where it remains for about 100 million years until it isuplifted and exposed again to continental erosion.
The life of atypical P atom is by human st~ndards indeed bleak.
It spends 99.9
percent of its time trapped in the sedimentary rocks of the earth.
Out of every 100,000,000 years it has only one 100,000 year stintin the ocean.
Since the volume of the surface water reservoir issmall compared to tha t of the deep wa ter reservoir, the P a tom
spends most of its ocean stints in the cold dark abyss.
Out ofevery 1000 years spent in the sea only a few years are spent in inwarm surface water!**The volume of the warm reservoir (plus one half of the main thermocline) "is only about one-tenth that of the cold reservoir (plus theother half of the thermocline). As water resides in the cold reservoir for 1000 years, it must reside in the warm reservoir about 100
years. Since the concentration of p in the warm reservoir is only 5
percent of that in the cold, the probability that P is removed byparticulate matter must exceed the probability that P is removed by
downwelling water by a factor of 20.
The removal time of P from thewarm reservoir must then be 20 times shorter than that of its companion water molecules.
Thus the residence time of a P atom in the warm
reservoir is only about 5 years.

20

Table 1-4. Example calculations for the constituents phosphorus, barium,and calcium.

I1

Phosphoruslf =\)rnix1 +\)river1= 0.01
1 + 30(3-0.15)

g = 1 -

1 +Csurrac1

CriverCriver1
1 + 30(0.3-0.1Y 0.14
Csurface\)river

Criver\)rnixCdeep\)river

Criver30 x 0.15= 0.95
1 + (30x3)

f x g =

1\)mix

Calcium

Cdeep

\)mix

1 +

1 -

Barium1 -1
1 + 30(25-24.8) = 0.14

30 x 0.1= 0.701 + (30xO.3)

1 -30 x 24.8= 0.011+(30x25)

Cdeep\)river Criver1
1 + (30x3)

T== 0.011
1 + C30xO.3) = 0.10

1
1 + 30x25 = 0.0013

Tmix
f x g

1000
0.01

1000
0.10

= 1 x 10 51000
50.0013 = 8 X 10

= 1 x 10 4.'

Table 1-5.

Surnrrary of model parameters for the elements P,

.

-._~.

Si, Ba. and Ca

category

ElEment

Csurface Cdeep
CriverCrivergffxg

Biolimiting

P

0.15

3.0

0.95

0.01

0.01

1 x 10 5

Si

0.02

0.7

0.97

0.05

0.05

2 x 10 4

ill

0.1

0.3

0.70

0.14

0.10

1 x 10 4

ca

24.8

25

0.01

0.14

0.00138 x 10 5

Biointermediate

21

T

yrs

If we are given the values for f and g for a particular constituent, we can determine the ratio of its concentration in surface water to that in river water, in deep water to that in riverwater, and in surface water to that in deep water. To do this wemust solve our equations for the concentration ratios in ter~ of
f, g, vriver' and vrnix.
The resulting equations are:- gvriverCriver

fg

vrnix

Cdeep
Criver

1 - fgfg

vrnix

Csurface

1Cdeep

1 - fgCsurface1

Vri ver

- g

1-13
1-14
1-15

Sample calculations for phosphorus, barium and calcium areshown in table 1-4 and a summary of the values of f and g obtained
for these three constituents and for silicon are given in table
1-5.

THE DISTRIBUTION OF BIOINTERMEDIATE CONSTITUENTS
Biointermediate elements merit a bit more discussion.
If
the composition of marine organic material is nearly constant thenthe surface water residuals for these elements should be nearly
the same from place to place in the warm surface ocean.
As can beseen in figure 1-6 this is to a good approximation the case forcarbon and for barium.
The variability from one place to anotherin surface waters between the lati tudes 40 0 N and 40°8 is small
compared to the difference between the concentrations of these
constituents in mean warm surface and mean deep water.
It thusappears that these constituents are taken up by plants in a cons tant proportion to phosphorus and when phosphorus has been depleted their concentrations cease to change.
One element, selenium, in the biointermediate category, has ageochemical complexity not shared by its brothers.
As shown in figure 1-7 Measures and his colleagues at MIT
found that selenium is partially depleted in surface waters.
Measures made separate determinations of total selenium and of
selenium in the +4 state (the concentration of +6 selenium could
then be determined by difference.)
As shown in figure 1-7 theconcentration of 8e+4 in surface water drops nearly to zero, whilethat of 8e+6 drops to 35% of its ambient deep water value.
Measures points out that the shape of the depth profile of 8e+4 resembles that of P0 4 and the shape of the Se+6 profile resembles thatof H 4 Si0 4 and infers that Se+4 is incorporated in organic tissue
and Se+6 in opal.
Like many oxidation-reduction reactions those for selenium
are kinetically impeded.
Measures points out that neither theratio of Se+4 to 3e+6 in surface water nor that in deep watercorresponds to the equilibrium thermodynamic ratio for the ambient
O 2 contents of these waters.
Hence he tentatively concludes that22

2 .3r----,..---r---r---~--...._-__.

,......
~:::::: 2.2
0

E
~
I

0

.....,

2.1

••

o

o

N

o

u 2.0
V'I

........
0"~-0..........

r--

60

WARM SURF. OCEAN ~
0

E

(1)
I

0

50

0

- 40
C

OJ

•

0

0

•
•
•
•• •
•
•
•
• • • /0
0

0

30
60

S

40

0

0

20
0
20
LATITUDE

40

60

N

Figure 1-6. Plots of the barium and total dissolved inorganic carbon (normalized to constant salinity) in surface waters as a function of latitude. The values for warm surface water (40 0 N to 40°8)
are remarkably constant.
These resul ts were obtained as part of
the GEOSECS program.
The barium measurements were made by J ahn
Edmond at M.l.T. and L.H. Chan at L.S.D. (375).
The total dissol ved inorganic carbon resul ts were obtained by shipboard ti tration (406).

23

-12

Se (10mol/kg)

1000

1500

500

0

~io

°1:1
•

I

E

~

.........

I

J-- 2

1'0.)
~

0....
·W

0

1

3~

4L

2000

2500 .1

.2

.3

.4

.5 .6

, x oX~ o~x~~ ( ; 0 0 0~••,-......

Se+ 4 /Se+ 6

••

0

0
0

x
x
x x

••
•

••
••
••
•
•
•
S +4 •••
e •••
•
•
•
•

0
0
0
0

XX

0
0
0
0
0

··1

x

0

+6 }

0
0

x
x x

l:Se

0

x •

0
0

0

x
x
x

0
0
0

•

0

~

•••
• ••
•
•
•
•
•

0

0

x
x

•

.•

0

x

Se

•

•

.

0
0

>O'x

x
x
x
x

• •
• •
•
••
••

0

•

•

.

.

.

.•

Figure 1-1. Plot of total selenium versus depth in the northeastern Pacific Ocean(29°N, 122°W).
Also shown are the concentrations of selenium in the two oxidation
states of this element (i.e. +4 and +6).
These measurements were made by Measuresand his coworkers at MIT (451).

some types of organisms must irreversibly transform Se+6 to Se+4
and others Se+4 to Se+6. Since about one third of the selenium inthe ocean is in the Se+4 oxidation state and two thirds in theSe+6 oxidation state the ocean-wide reduction probabili ty must be
about one half the ocean-wide oxidation probability.

ESTIMATION OF INPUT RATES
Thus far we have not mentioned how the estimates of the concentration of elements in average river water is determined.
As
it turns out these estimates are not so easily obtained.
There
are several reasons for this.
1) The reliability of many of the analyses of at least thetrace constituents in river waters is open to question.
2) Man's activities have changed the chemistry of manyri vers.
3)
The dissolved concentration of a constituent in riverwater may not be an adequate measure of the input of thisconstituent to the sea. Some constituents may be carriedin organic tissues which will decompose in the ocean.
Others will partially precipitate in the estuaries.
Others will desorb from silicate detritus in estuaries.
4)
Ri vers differ from one another in their chemis try.
Agiven river may have a different chemistry during flood
stages than during low water periods.
Thus averaging is
not easy.
Despite these difficulties attempts have been made to assess
the average input rate of a larger number of elements to the sea.
These estimates are summarized in table 1-6.
This table also
gives estimates of the mean concentrations of these elements in
sea water.
For those elements where both pieces of information
are available residence times (i.e., mean ocean concentration multiplied by ocean volume divided by mean river concentration multiplies by annual river flux) are also given.
The entries in thistable are taken from standard compilations and do not always agree
exactly with the ones employed in this chapter.
As none of the
input rates are known to better than 20% these minor differences
are not significant.
In our discussion and in table 1-6 the rate of input to thesea of the various constituents of seasalt is determined by measurements of their concentration in river water.
This approach is
not always reliable.
Estimates can also be made from the rate ofaccumulati~n of the constituent on the sea floor.
Phosphorus provides a good example. The historical data on phosphate in rivers
is sketchy and often inaccurate.
Present day rivers carry excess
phosphate released from agricul tural land and from sewage plants
and some of the phosphorus carried by rivers is a constituent partof living organisms and their residues and hence is not included
in analyses of soluble phosphate.
Thus estimates of the steadystate rate of phosphorus addition based on river data are very uncertain.
An alternate estimate can be made by considering therate at which phosphorus has been removed from the sea into marine
sediments during the last several thousand years. This is done bymeasuring the accumulation rate and phosphorus content of various

25

Table 1-6. Element concentrations in average river water and in averageocean water. Also given are the corresponding mean oceanic residence times.

Atomic'No.

Element

1
2
3
4
5
6

H (as H 2 O)
He

7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
30
31
32
33

34
35
36
37
38
39

40
41
42

11

Be
B
C (inorganic)(organic)
N (dissolved N2 )(N03 )
0 (as H2 O)(dissolved 02)
F

Ne
Na
Mg

Al
Si
P

S
Cl
Ar
K
Ca
Sc
Ti
V

Cr
Mn

Fe
Co
Ni
Cu
Zn
Ga
Ge
As
Se
Br
Kr
Rb
Sr
y

Zr
Nb

Mo

Cone. ~ean River 1(10- moles/kg)

1.7
1.7
-

*5.3

2.2 x 10 2
1.6 x 10 2
1.9
1.9 x 10 2
1.3-

3.4 x 10 21
3.6 x 108.9 x 10- 51
2.1 x 102.0 x 10- 2
1.9 x 10- 21
1.5 x 10- 1
7.2 x 103.4 x 10- 3
3.8 x 10- 2
1.6 x 10- 11
4.6 x 101.3 x 10- 3-

2.3 x 10- 2*2.5 x 10- 31
2.5 x 10--

1.8 x 10- 21
6.9 x 10*7.9 x 10- 3-

5.2 x 10- 3

26

Conc~

(10-

Mean Sea 2moles/kg)

5.4 x 10 7
1.8 x 10- 3
2.5 x 10 1
(6.5 x 10- 5 )4.2 x 10 2
2.3 x 10 34
5.8 x 10 2
3.0 x 10 1
5.4 x 10 7
2.2 x 10 2
6.8 x 10 1
7.5 x 10- 3
4.7 X 1054
5.3 X 10(3 x 10-2 2 )1.0 X 10
2.3
2.8 X 1045.5 x 1051
1.5 x 1010.2 X 10 33
10.3 X 10 5
(1.5 X 10- )«2.0 x 10- 22 )2.3 X 10- 3
4 X 105 x 10- 33(1 X 10- )(3 X 10- 5 )8 X 10- 3
4 X 10- 3
6 X 10- 3(3 X 10- 45 )7 x 10- 2
2.3 X 101.7 X 10-2 3
8.4 X 10
3.4 x 10- 3
1.4
8.7 x 10 1 4(1.5 X 10- )(3 x 10- 4 )5«5 x 10- 1 )1.1 X 10-

T

3

(yrs)

- 5
5.7 X 10
9.6 X 10 6
- 105

5.0 x

8.3 X 10 77
1.3 X 10
6.2 x 10 2
2.0 X 10 4
6.9 x 104--

1.2 X 10 67
1.1 X 10
3.7 X 10 43
4.5 X 10 3
8.2 X 10
1.3 x 10 31
5.4 x 10
3.4 X 10 2
8.2 X 10 3
9.7 X 10 22
5.1 x 109.0 x 10 3-

3.9 X 104-4
2.6 X 10
1.3 x 10 8-

3.0 x 10 66
5.1 X 10
7.4 X 10 2-

8.2 X 105

Table 1-6. (continued)
AtomicElement

No.

43
44
45
46
47
48
49
50
51
52
53
54
55
56
57
58
59
60
61
62
63
64
65
66
67
68
69
70
71
72
73
74
75
76
77
78
79
80
81
82
83
90
92

Tc

Ru
Rh

Pd
Ag
Cd
In

Sn
Sb
Te

I
Xe
Cs
Ba
La

c:e

Pr

Nd
Pm

Sm
Eu

Gd
Tb
Dy

Ho
E~

TIn

Yb
Lu
Hi'
Ta
W

Re
as
I~

Pt
Au
Hg
Tl

PbBi
Th

U

Cone. Mean Rive~1(10- 6moles/kg)-

2.8 x 10- 3
8.2 x 10- 3*5 x 10- 24

2.6 x 104.4 x 10- 1
3.6 x 10- 4
5.7 x 10- 45
5.0 x 102.8 x 10- 4-

55.3 x 106.6 x 10- 6
5.1 x 10- 5
6.3 x 10- 6*3.0 x 10- 4
6.1 x 10- 6
2.4 x 10- 5
5.9 x 10- 6
2.3 x 10- 5
5.7 x 10- 6*1.6 x 10- 45

1.0 x 104
*3.5 X 104.8 X 10- 3--1 x 10- 3

Cone. Mean Sea 2(10- 6 moles/kg)--(2.5 x 10- 45)
7 x 10(1 X 10- 6 )(4 x 10- 6 )1.2 x 10- 3

4.4" x 10- 1
5.0 x 10- 4
2.2 X 10- 13
1.0 X 103 X 10- 55
2 x 104 x 10- 56
2 x 10-

1

3

(y~s)

- 2
3.5 X 10- 3
5.7 x 10
3.4 X 105-

3.3 x 10 53
8.8 X 10 3
3.2 x 10
1.4 x 10 3
3.1 x 10 3
2.8 X 10 3-

4 X 10- 6
7
9 X 10- 6
6 X 10- 7
9 X 106 x 10- 6
2 X 10- 66
5 X 10- 7
8 x 105 X 10- 67
9 X 10- 5«4 x 10- )«1.4 x 10- 5 )6 X 10- 45(2 x 10- )2.9 X 10 33
5.3 X 10
4.6 X 10 33
5.6 X 10
7.8 X 10 24
1.3 X 10
8.1 X 10 33
5.3 X 10
8.5 X 10 33
6.2 X 10(2.5 x 10- 65 )(5 x 10- 5 )6 x 101 x 10- 54(1 x 10- )6«3 x 10- 2 )1.3 X 10-

9.7 X 10 2

-

-

-

4

5.6 x 10-8.1 x 10 1
-5 x 10 5lr~in sou~ce:

A summa~y by Martin and Meybeck (631) with values
for F, Se, Y, I Dy, W, and Mg, values with * are taken from Tu~ekian(628).
source: A summary by Bruland (456). n1e values in parenthesis
are highly uncertain.
3The residence time is obtained by diViding the total amount of the element dissolved in the sea by the amount delivered by rivers each year.

2Main

27

types of marine sediments (s ee table 1-7).
If the input andoutput rates are equal, the mean phosphorus content of river watercan .then be estimated.
Froelich, who has carried out the mostdetalled analysis of this sort, concludes that 3.5 x 10 15 moles ofphosphorus are dissolved in the sea and that its sedimentationrat~ is about 3.5 x 10 1 0 moles of phosphorus per year.
The
resldence time obtained in this way is 100,000 years (the value
used in our discussion of phosphorus).
It is of interest to notethat. whil~ 80% o~ the phosphorus leaves the ocean wi th biologi cremalns (In keeplng with our first order model) half of thisphosphorus is bound in CaC0 3 rather than in organic tissue.
A similar problem exists with the river based estimatesfor carbon.
Much of the bicarbonate ion dissolved in river wateroriginates from atmospheric C02 which neutralizes the bases
produced during rock weathering processes.
This CO 2 is recycled
ocean-atmosphere carbon rather than new carbon being added byweathering or volcanism.
While means are available for correcting
for the contribution of recycled carbon they are not entirely
adequate.
Again sediment accumulation rates of carbon can be used
to provide an independent cross-check.
Using this approach
Froelich obtains a carbon residence time of 180,000 years (634).

HORIZONTAL SEGREGATION OF CONSTITUENTS IN THE DEEP SEA
The simple two box model of ocean chemistry has permitted us
to describe in a quantitative manner the primary features of themarine chemistry for biologically utilized consti tuents.
In thereal world the concentration of a given constituent shows considerable geographic and depth variabili ty in the cold ocean.
Asshown in figures 1-1 and 1-2 the concentrations of the biologically active constituents increase with depth in the North Pacific toa mid-depth maximum and then decrease somewhat toward the bottom.
This type of profile characterizes the Pacific and Indian Oceans(to abou t 40° S) .
In the Antarctic Ocean the profiles show muchless structure.
In the Atlantic the concentrations of these constituents increase all the way to the bottom.
Figures 1-8, 1-9, 1-10 and 1-11 show the geographic patternsof the concentrations of H4 Si0 4 , Ba++, N0 3 - and O2 along the 4000
meter depth horizon. For several reasons these maps do not have afamiliar look.
First, were the oceans drained to a depth of 4000
meters, their geography would look quite different. The mid-ocean
ridges would rise above the surface and would divide the Atlantic
into two separate oceans, the Indian Ocean would be divided into a
series of basins and the Pacific into a major and a minor basin.
The continents would be larger and the oceans smaller.
Hence theshapes of the margins of the 4000-meter horizon do not closely resemble the margins of the present ocean.
Further, in order tomake these maps page size, we have stretched the distance between
la ti tude lines by a factor of two.
Final'ly, we have not compensated for the poleward shortening of the distance around the globe
on a given latitude.
This makes the polar oceans look too large
and the equatorial oceans too small.
As can be seen on these maps, the concentration of the nutrient constituents increase "down" the Atlantic and "up" the Pacific28Table 1-7. Summary of phosphorus burial rates in marine sediments
as obtained by Froelich and his coworkers (634).

Form

Burial rate
10- 9

moles
cm 2 yr

Organic matter
CaC03
Hydrothermal
Phosphorite
Fish debris

<0.2

TOTAL

---10

4
4
1
<1

Burial rate
10 9 moles
yr

14
14
4
<4
<1
-""35

Table 1-8.
Horizontal segregation of biologically utilized constituents between the deep North Atlantic and the deep North
Pacific.
To permit meaningful calculations for the biointermediate elements, carbon and barium, only the deep water excess is
considered.
Note that the horizontal segregation is more pronounced for those elements incorporated into hard parts (silicon
and barium) than for those incorporated into soft parts (ni trate
and phosph.a te) .
Constituent.

(Cdeep - Csurface)pacific(Cdeep - Csurface)Atlantic

N03

2

P0 4

2

H4 SiO 4

5

Ba

4

29

o
20

40

60

60

20

20

Figure 1-8. Distribution of dissolved silicate at 4000 meters
depth in the world's major ocean basins.
The boundaries of theocean are those at which the 4000 meter horizon abuts the sediment.
The Atlantic Ocean basins are on the left, the Indian Ocean
basins in the center and the Pacific 0 cean is on the right. Basedon results obtained during the GEOSECS program (424, 425, 426).

r--..---~--or---..,...---.r--..,..;.-.....,.--.n-;..-~~60

40

Figure 1-9. Distribution of dissolved barium at a depth of 4000.meters in the world's major ocean basins.
Based on results obtalned
by John Edmond of M.I.T. and
L.H. Chan of L.S.U. on GEOSECS samples (375).

30

Figure 1-10. Distribution of dissolved nitrate at a depth of 4000

meters in the world's major ocean basins.
Based on resul ts obtained during the GEOSECS program (424, 425, 426).

Figure 1-11. The distribution of dissolved oxygen of 4000 meters
depth in the world's major ocean basins.
Based on results obtained
during the GEOSECS program (424, 425, 426).

31and Indian Oceans.
Dissolved oxygen has, as expected, the reverse
sequence of concentrations.
Like the vertical distributions
thehorizontal distributions of these properties have similar ~jor
structures.
However, the details differ from property to property.
For example, oxygen shows a strong maximum along the margin of the Antarctic continent.
Equi valent minima are not seen
for nitrate or silicate.
As shown in table 1-8, the Pacific toAtlantic enrichment of the biologically utilized constituents isdifferent for the hard-part constituents than for the soft-partconstituents. *The difference between the enrichment of these biologi cally
active constituents in the deep Pacific and in the deep Atlantic
depends on the deep water circulation pattern in the world ocean.
About half the water being fed into the deep ocean is generated at
the northern end of the Atlantic Ocean.
As shown in figure 1-12
this water flows down the western Atlantic. It enters the eastern
Atlantic through a break in the Mid-Atlantic Ridge located close
to the equator.
It enters the Antarctic through the southern endof the western basin of the South Atlantic and joins the eastward
floWing circumpolar current.
Part of the water in this circumpolar current penetrates northward into the basins of the Pacific;
part passes through the Drake passage and starts a second triparound the Antarctic continent.
There is one major complication:
the circumpolar flow isjoined by water which has been recooled at the surface of the
Antarctic Ocean.
This cold water sinks along the margin of the
Antarctic adjacent to the South Atlantic.
Thus it is a mixture ofthis recooled water and water emerging from the Atlantic thatflows into the Indian Ocean and into the Pacific Ocean.
As weshall see, the water formed in the Weddell Sea area of the Antarctic continental margin is virtually unmodified in its nutrient
constituent concentration.
Oceanographers tell us that the production of new deep wateris balanced by the upwelling of water to the surface throughout
the world ocean.
Wi th this in mind, it is not difficu1 t to see
why there is a tendency for a biologically acti ve element to beenriched in deep Pacific water relative to deep Atlantic water.
Plant acti vi ty has partially depleted these nutrients in surface
water sinking into the deep Atlantic.
As this water passes
through the Atlantic, particles bearing these biolimiting cons ti tuents fall into it from the surface.
Its nu trient content isthereby increased. This enrichment process continues as the water
moves into the Indian and Pacific deeps.
As it proceeds along its
path (see figure 1-13), the current is continually weakened by upwelling.
The water lost in this way returns via surface currents
to the deep water source region (the Norwegian Sea).
The nutrients, however, are extracted by organisms and returned to theabyss in particulate form.
*These enrichment factors are calculated by subtracting the surface
water concentration for a given constituent from its deep water
concentration and dividing this difference for the North Pacific bythat for the western North Atlantic.

32

40

20

w

o

:::>

ff-

<:t

.....J

40

60 ......------a.~~-.a.--

....

60

~~~-----~;a,.....

100

140

......

~-----60--01........

LONG I TUDE

Figure 1-12. Flow pattern at a depth of 4000 meters.
The major
puts to this horizon are North Atlantic Deep Water (NADW) whichters at the northern end of the western basin of the Atlantic
Weddell Sea Bottom Water (WSBW) which enters from the margin of
Antarctic continent adjacent to the South Atlantic.

33

inenandtheEUROPE-ASIA

AFRICAANTARCTICA

ATLANTIC

INDIANPACIFIC

Figure 1-13. The upper diagram is an idealized map of the patternsof de e p wa t e r flow ( sol i d 1 i n e s ) and surf ace wa t e r flow (d ash e dlines) .
The large circles designate the sinking of NADW in the
Norwegian Sea and the recooling of water along the perimeter ofthe Antarctic Continent; the shaded circles indicate the distributed upwelling which balances this deep water generation.
The
lower diagram is an idealized vertical section running from theNorth Atlantic to the North Pacific showing the major advective
flow pattern (solid lines) and the rain of particles (wavy lines).

34

The net result is a steady push of nutrients toward the deep
Pacific.
They can ride the deep current away from the Atlantic
bu t not the returning surface current.
This creates an increasein nutrient content along the path of the deep current. The gradient has become sufficiently steep so that horizontal mixing at
depth moves the nutrients back toward the Atlantic as fast as they
can be exported by the deep current.
Thus a dynamic balance isachieved which stabliizes the Pacific-ward concentration gradient.
It is the interac tion between oceanic circula tion and biologic
ac tivi ty that genera tes the enrichment of these elements in thewaters of the deep Pacific relative to those of the deep Atlantic.
An analogy will clarify the essentials of this cycle.
The
exhibi ts in a fun house are loca ted on two levels.
The upper
floor has a large conveyor belt that moves from right to left; thelower floor, a belt that moves from left to right.
Those who enter are free to observe the horrors in any order they wish. There
are innumerable escala tors from the lower to the upper level.
However, there is only one escalator from the upper to the lower
level, located at the end of the upper belt.
Those who venture to
the upper level are harassed by mons ters lurking in dark alcoves.
These monsters grab the unsuspecting visitors and, after a suitable frightening, drop them through holes to the lower level. The
average fun-seeker has to ride to the upper level many times toview all its mysteries before leaving the fun house.
If, on a busy Sunday afternoon, we were to snap on the lights
suddenly and photograph the distribution of people, we would findmany more fun-seekers downstairs than up (the untiring monsters
qUickly track down those who step from the escala tor tops andpromptly throw them back down to the lower level, relatively close
to their point of ascent).
On the lower floor, more people would
be at one end of the building than at the other (the steady movement of the lower-level belt would carry the wanderers toward thecrowded end).
As long as people remained in the fun house for aperiod covering several belt cycles, their distribution would notdepend on where they entered or left.
Ra ther it would depend on
the speed of the belts and the efficiency of the monsters.
Obviously the fun-seeker is the limi ting nutrient and themonster is the plant.
The belts and escalators represent the organized flow of wa ter, and the wandering of the people is theturbulent mixing superimposed on this organized flow.
Al thoughnot perfect, this analogy does capture the important factors
influencing the distribution of nutrients in the sea.
Now let us consider the more detailed fea tures seen in thevertical profiles in figures 1-1 and 1-2 and in the horizontal
distribu tions of properties at 4000 meters in figures 1-8, 1-9,
1-10, and 1-11. The mid-depth nutrient element maxima seen in the
Pacific Ocean are genera ted by the general upwelling of bottom
water combined with the progressive destruction of the particulatematter which falls from the surface into this rising column.
Thisfeature is absent in the Atlantic because of a strong along-bottom
counterflow of water laden with nutrients from the Antarctic into
the Atlantic.
This along-bottom counterflow is possible because
the deep water formed in the Weddell Sea is more dense than thatformed at the northern end of the Atlantic.
The fact that the35

WESTERNATLANTIC
H4 Si0 4E~

fLmol/kg

I

to..
W
o

4

5'---_.a.-_.&-.IL.-...a-

60
S

50

40

30

_ __....._~

20._.a..._~

10

a

10

LATITUDE

20_ a . _....._

30

40

_..I

50

~

60
N

Figure 1-14. Vertical section along the western basin of the Atlantic Ocean showing the distribution of dissolved silicate. Contours
are drawn at intervals of 10 micromoles per kilogram. The main features to be seen are the southward penetrating tongue of low silicate North Atlantic Deep Water, the northward penetrating wedge ofhigh silicate Antarctic Bottom Water, the northward penetrating
tongue of high silicate Circumpolar Intermediate Water, and the low
silicate content of surface water everywhere except in the Antarctic.
Two horizons of constant density are shown by the dottedlines (the upper is the 00 = 27.0%0 isopycnal and the lower the04
45.2%0 isopycnal).
In the Antarctic the concentration
contours neatly parallel the densi ty horizons.
Based on resul ts
obtained during the GEOSECS program (424).
The tracks along which
this and subsequent sections are drawn can be seen on the foldout
map.

36

I

I--

a..
w

o

4

51.70
S

. l - _ - I - _........_ . - . . . . . _......._.....&......a.-......L._---L_---II~_I.II..___..........

60

50

40

30

20

10

0

LATITUDE

10

20

30

40

50

__

60
N

1-15. Vertical section along the western side of the Pacific Ocean showing the distribution of dissolved silicate. Contours
are drawn at 10 micromoles per kilogram intervals. As for the Indian Ocean. (see figure 1-17) the silicate content of deep water
increases toward the north.
Again a trough of low silicate water
centered between 40 and 50 0 S is seen.
Two horizons of constantpotential density are shown by the dotted lines (the upper is the00
27.0 0 /00 isopycnal and the lower is the 04 = 45.2 0 /00
i sopycnal) .
Based on resul ts obtained during the GEOSEC.S program

F~gure(425).

37

'E2

-2
E

~

~

I

J0-

n.

W
0

3
flO
120

4

~O
S

WESTERN
INDIAN
H4 Si0 4
fLmol/kg

4

50

40

30

LATITUDE

20
S

5

60

50

S

40

30

20

LATITUDE

10

0

10

20
N

Figure 1-16. Vertical section of dissolved silicate concentrationin the eastern Atlantic to the south of the Rio Grande Rise whichisolates the deep eastern basin of the central Atlantic from theAntarctic Ocean.
The low-silicate deep water on the right is theextension of the North Atlantic Deep Water tongue.
It lies against the high - silicate circumpolar deep water which hugs theflank of the Antarctic continent. Again a horizon of constant potential density is shown by the dotted line (the 04 = 45.2 0 /00
isopycnal).
Based on results obtained during the GE03ECS program(424).
Figure 1-17. Vertical section along the western basins of the Indian Ocean showing the distribution of dissolved silicate.
Contours are drawn at 10 micromoles per kilogram intervals. By contrast to the sections for the Atlantic shown in figures 1-14 and1-16 silicate concentrations in deep water increase northward of
40°3.
A feature of interest is the trough of low silica content
water between the la ti tudes of 30 and 50° S.
It represents thecontinuation of the North Atlantic Deep Water tongue.
Two horizons of constant potential density are shown by the dotted lines(the upper is the 00 = 27.0 0 /00 isopycna1 and the lower is the04 =45.2 0 /00 isopycnal).
Based on results obtained during the
GEOSECS program (426).

38

<:)

0

SOUTH ATLANTICo SOUTH INDIAN00~

SOUTH PACIFIC

0

C"

,

0

0

a

~

E

0

::l.

v-

a

v-

0 0

0

--en
J:

~

00
~

30...--......-..,..-......--......--..,...-.....- ......----.
o SOUTH ATLANTICo SOUTH INDIAN~o

SOUTH PACl FIC

8
<:)

~

o
z

o
8

0.....-------'---------M ~aJ..~~0 ~
60
50
40
30
LATITUDE oS

Figure 1-18. Dissolved silicate and nitrate content of surface
water as a function of latitude in the southern oceans.
Both constituents are found in low concenration in temperate surface
waters and in high concentrations in Antarctic surface waters.
The transi tion from the high to low values occurs at higher latitUdes for silicate than for nitrate presumably reflecting the predominance of silicate utilizing plants (i. e., diatoms) in polar
surface waters.
Based on results obtained during the GEOSECS program (424, 425, 426).

39

mid-depth extrema are deeper for consti tuents carried in hardparts (i.e., H4Si04, alkalinity and barium) than in those carriedin soft parts (i.e., N0 3 and P0 4 ) is the result of the fact thatmineral solution occurs on the average at greater depths than respiration.
This same factor accounts for the greater Pacific-ward
enrichment of the products of hard part solution than for the products of respiration.
The north-sou th sections of silica concentration for the three oceans shown in figures 1-14, 1-15, 1-16 and1-17 nicely portray these features.
The high oxygen content of deep water found in the Antarctic
adjacent to the Weddell Sea is a consequence of the input of new
bottom water in this area.
The absence of corresponding anomalies
in N03 and H4 Si0 4 is, however, unexpected.
The new deep water
formed during the dark winter in the Weddell Sea is in contact
with the atmosphere long enough to be cooled and long enough fordissolved oxygen to be replenished but does not lose a significant
portion of its nutrient elements to plants.
Figure 1-18 (and also figure 1-6) clarify the distinction between cold and warm surface wa ters by showing that the nu trient
content of high lati tude surface water is more akin to that ofdeep water than to that of warm surface water.
The nitrate values
in surface water are all nearly zero until 40 0 S is reached.
Then,over the latitude range 40°8 to 55°S, they rise to the rather highvalues characteristic of polar waters. The silicate concentrationis also low in temperate surface waters.
These low silicate values of temperate surface waters extend to higher latitudes than dothose for nitrate.
This is consistent with the fact that diatoms
dominate the plant communities found at high latitudes.
SUMMARY

In this chapter, two important aspects of marine chemist ry
have been discussed.
First we have examined the factors controlling the distribution of the biologically utilized constituents
within the sea.
Because these species remain in the sea for times
far longer than that required for the ocean to miX, their distributions do not reflect their points of entry to the sea or theirpoints of removal from the sea.
Rather their distribution reflects the interaction between the large-scale circulation patternof the sea and the life cycle in the sea.
The latter drives these
consti tuents from surface to deep water.
The former pushes themfrom the deep Atlantic around toward the deep northern Pacific.
The differences in the vertical and horizontal distributions ofthese constituents are the result of differences in the depth towhich the various particulate phases fall before being returned tosolution. Some of these species are totally depleted in warm surface water; others are only partly depleted.
The degree of depletion depends on the relationship between the biochemical demand
and the ocean availability of the particular constituent.
Second we have examined the factors controlling the residence
time of biologically important constituents in the ocean.
We haveshown that both the extent to which a consti tuent is depleted inwarm surface water and the extent to which the phases so produced40

are able to survi ve destruction in the deep sea influence
elapsed time between entry into the sea and burial in
sediments.

the
the

PROBLEMS

1-1 Constituent x has an average concentration of 2 milligrams per
Ii ter (mg/l) in ri ver wa ter, 4 mg/l in surface wa ter and 16
mg/l in deep sea water.
What are the values of g,
and T
for this constituent?

-r,

1-2 For constituent y 20% of the material falling to the deep sea
is preserved in the sediment (i.e., 80% is recycled).
Fifty
percent of this constituent reaching the warm ocean is fixed
into particulate matter which falls back into the deep sea.
Its concentration in average river water is 10 micrograms per
Ii ter (~g/l). What is its average concentration in deep sea
water?
In surface sea water?
What is its average residence
time in the sea?

1-3 Marine organisms use Sr as well as Ca to build their CaC 03
shells.
The ratio of Sr to Ca in shells is only one-fifththat in the water in which they dwell.
If this is the onlymeans by which Sr is removed from sea wa ter, how would theSr/Ca ratio in ri ver water compare to that in average sea
water?
What would be the ratio of the residence times of Srand of Ca?
1-4 If, at some time in the past, deep water formed at the northend of the Pacific rather than at the north end of theAtlantic, how would the chemistry of the sea differ from what
it is today?

1-5 I f organisms were to become less efficient than they now arein consuming the organic debris falling to the deep sea, howwould the chemistry of the sea change?
On what time scale
would these changes occur?
1-6 Were life in the ocean suddenly to be destroyed, what changes
would occur in the nutrient cons ti tuent chemis try of surface
water on a time scale of a hundred of years?
What changes
would occur on the time scale of 100,000 years?
SUPERPROBLEM #1

A biologist named Johannes Theft constructs a special fourreservoir aquarium, designed to portray the distribution of nutrients in the world ocean (see figure 1-19).
The upper two reservoirs are open to the air and illuminated by aritifical lights.
The lower two are sealed from contact wi th the air and are dark.
Stirring devices are installed so that the waters in each reservoir remain well mixed. Pumps are installed to transfer water between the reservoirs at fixed ra tes.
Pumps are installed to add

41

I

I
8--

o

C

INDOPACATLANT IS

-

T

10

....2

/

I

I

VOLUMES

RAe: 8 LITERS/HR
ReA: 2 LITERS/HR

VA : 600 LITERS

Va: 300 LITERS
Vc : 3000 LITERS
Vo : 6000 LITERSPUMPING RATESINA = I L ITER/HR
aUTo: I LITER/HR

Rec = 10 LITERS/HR
RCB : 4 LITERS/HR
Reo = 8 LITERS/HR
Roc = 2 LITERS/HRROA = 5 LITERS/HR
RAD : 0 LITE RS/HR

Figure 1-19. The design of Johannes Theft T s four-reservoir aquarium of superproblem #1.

42

new sea wa ter to reservoir A and remove an identical amount ofwater from reservoir D.
Theft fitls the system with sea waterwith an N03 concentration of 30xlO- moles/liter and a P04 concen6
tration of 2xlO- moles/liter.
The input water has the same concentration of these nutrients.
Theft inoculates the surface water
wi th algae and the deep wa ter wi th bac teria.
Theft adds a dozen
or so snails to each deep reservoir.
Finally, Theft ins talls afiltering system in each surface reservoir, so that the water canbe purged of living plants and suspended particulates.
These filters are transferred at frequent intervals to the underlying reservoirs and "back-flushed", dispersing their catch in the deep
wa ter.
Theft views his creation with some satisfaction and then puts
it into action.
The lights go on, the pumps go on, the filteringsystems go on, and the mixers go on.
He then begins a program
monitoring the N0 3 and P04 content of the four reservoirs.
As expected, the concentrations in the surface reservoirs quickly drop
to near zero values.
That in the smaller of the two deep reservoirs, which is appropriately named Atlantis, also begins to drop
but levels out at a constant value.
After some time, all the concentrations stabilize; steady state has been achieved.
Theftnotes with pride that no organic slime accumulates on the sides orthe bottom of the deep reservoir.
The bacteria and snails dotheir job and consume the material back flushed from the filters.
What are the steady state N03 and P0 4 concentrations in Atlantis? In Indopac? If the dissolved oxygen content of the water
descending from the surface is 350xlO- 6 moles/li ter and, if the
C/N/P ratio in the material caught on the fil ters is 105/15/1,
then what is the 02 content of the waters in Atlantis and in Indopac? Roughly how long did Theft have to wait for steady state tobe achieved?
After showing off to his colleagues
and students the success of his crea tion, Theft becomes a bi t bored and decides tochange some things.
First he halves the NO 3 and PO 4 content of
the feed water.
As he expects, after a suitable time the concentra tions of nu trients drop to and remain at half their previous
steady state values.
The productivity of the plants also falls tohalf its previous value.
After restoring things to their original state, Theft cuts
off three of the pumps and merely circulates water at the rate of
10 li ters per day from tanks A to tank B to tank C to tank D andback to tank A. He is amazed by what happens to the nutrient concentration-and snails in Atlantis.
Calculate the new steady state
nutrient concentrations in both Atlantis and Indopac.
What fatebefell the snails in Atlantis (but not those in Indopac)?
Once again Theft restores the tanks to their original mode ofoperation.
While contemplating what experiment he might next try,a geochemical colleague named Spencer Crane drops in.
After viewing the aquarium for some time he tells Theft that, while his system is a very clever one, which does indeed reproduce the major
nutrient distrubution in the ocean, it has one flaw.
"Your Indopac reservoir always seems to end up wi th a nu trient concentration equal to that in the feed water," says Crane.
"That is hardly the si tua tion in the ocean." As a means of rec ti-

43fying this deficiency, Crane suggests that the outflow be switchedfrom reservoir D to reservoir B, and that a second set of muchsmaller filters be run in parallel with the existing filters, andthat their catch be discarded rather than back-flushed into theunderlying reservoirs.
Theft sees the point and makes thechanges. He first drops the pumping rate D to B, C to D, and D to
A by one Ii ter per hour so that water balance is maintained forthe new arrangement.
He sets the relative catches of the primary
filter and secondary filter (i.e., added throwaway) to be 100.
In
this case, what would be the steady state N03 and P0 4 concentrations in Atlantis?
In Indopac? What would the O2 concentrations
be? If the area of the surface of reservoir A is 3 times that ofreservoir B, what would be the ratio of the plant debris rain
rates (per unit surface area) from these reservoirs?

NOTE: On page 570 are listed some constants that will prove handy
for the problems in this and in subsequent chapters.
LAST MINUTE ADDITION (SEPTEMBER 1982)
As this book was being written and corrected exciting results
were being obtained on the oceanic distributions of Pb, Be (722),
and Ge (728), and the rare earths.
The list of classifiable elements grows rapidly and the complexity of the oceanic distribution
patterns become more complex.
We shall have more to say aboutsome of these findings in chapter 4.

44

Ckaptelb 2
THE SEDIMENTARY SINK
FACTORS INFLUENCING THE DISTRIBUTION OF SEDIMENTARY CONSTITUENTS

INTRODUCTION
On the average the opal and calcite generated by the organisms living in the sea make up about 50% of the sediment accumulating on the sea floor.
The geographic distributions of thesephases are, however, far from uniform.
Sediments consisting ofnearly pure opal and of nearly pure calcite cover large reaches ofthe sea floor.
So also do sediments nearly free of both of thesephases.
The abundance of these phases at any point on the sea
floor reflects both their rate of production in the overlying
water column and the extent to which they are subjected to solution on the sea floor.
Since both the rate of production and theextent of solution of opal and calcite are related to distributionof nutrient constituents in the sea the abundance patterns ofthese phases in sediments offer valuable clues regarding the operation of the earth's largest chemical plant.
Our major goal inthis chapter will be to exploit this information.

SEDIMENT TYPES
Deep sea sediments consist of three major components.
The
mos t voluminous of these is detri tal rna terial deri ved from themechanical and chemical fragmentation of the materials which makeup the continents (see table 2-1).
Most of this material is in
the form of alumino-silicate minerals.
It is dropped by winds onto the sea surface and is disgorged by rivers into coastal waters.
As might be expected, this detritus accumulates most rapidly alongcontinental margins and more slowly in areas of the sea floor more
remote from land.
Sediment cores taken more than a few hundred
kilometers from land masses show detri tus accumulation rates between 0.2 and 1.0 grams per square centimeter of ocean floor per1000 years.
Deep sea sediments rich in this component are called
red clays.
So-called authigenic minerals in sediments are formed byspontaneous crystallization within the sediment or water column.
They make up only a small fraction of the total sediment.
Themost important of these is the iron-manganese oxide material in
sediments formed through reduction of these metals at depth in thesediment column coupled with upward migration and redeposition in
the oxygenated upper layers of sediment. This material is also a
by-product of hydrothermal acti vi ty near the ridge cres ts.
Hot

45

Table 2-1. Mineralogy of alumino-silicate detritus transported
from the continents into the sea via rivers and air.

MineralChemical Composition

QuartzOrthoclase
Plagioclase
Kaolinite
IlliteSmectite
Chlori te
Vermiculite

Si02
KAISi30S
xNaAlSi30S + (l-x)CaA12Si20S
A12Si20S(OH)4
KAl3 Si 3010 (OH) 2
AI2Si4010(OH)2.xH20(Mg,Fe,AI)6(Al,Si)4 0 10(OH)8
Mg3S i 4010(OH)2· xH 20......

-z
~
0

0

OJ

a:
<t
U

u

-

2

Z

•
•
• •

<t

(!)

a:
0

.01

•

•

4

•
•

•

•

100
10
0.1
I
3
2
SEDI MENTATION RATE (g/cm ,0 yr)

Organic carbon content of marine sediments as a function of bulk sediment accumulation rate. Results summarized by
Ross Heath and his colleagues (531).
Figure 2-1.

46

waters entering the sea in these regions carry with them sizeable
amounts of reduced Mn and Fe.
On venting, these metals encounterthe dissolved oxygen in the deep sea water, become oXidized, flocculate, and settle to the sea floor.
In some ridge crest sediments, these oxides of Fe and rtin make up several tens of percentof the sed iment.
The biogenic component of sediment consists mainly of two
substances that are produced as hard parts of organisms.
These
are calcite and opal precipitated by plankton living in the watersof the surface ocean.
As mentioned in chapter 1, calcite isformed by coccoli ths (plants) and by foraminifera (animals); opal
is formed largely by diatoms (plants) and radiolarians (animals).
The distribution of calcite and of opal in ocean sediments is not
at all uniform.
In some places on the floor of the Antarctic
Ocean we find an enormous amount of opal mixed with the ever present detrital alumino-silicates.
On the crests and upper flanksof mid-ocean ridges we find that the ever-present blanket of clay
is diluted with a large amount of calcite.
At great depths in thetemperate ocean, we find neither opal nor calci te; alumino-silicate debris accumulates virtually alone.
Soft tissue residues rarely make an important contribution to
deep sea sediments (a tribute to the efficiency of the animals andbacteria that feed on it). We do not have a good understanding of
the details of the accumula tion pa ttern for this rna terial.
The
results in hand do, however, show a correlation between increasing
sediment accumulation rate and increasing organic carbon content(see figure 2-1).
Sediments that accumulate at rates less than
one centimeter per thousand years have an organic carbon content
of only a few tenths of a percent.
By contrast, where sediments
accumulate at rates greater than 10 centimeters per thousand
years, organic carbon contents of several percent are found.
Onlyin anaerobic basins such as the Black Sea and Carioca Trench are
sediments containing more than 10 percent organic carbon found.

DISTRIBUTION OF OPAL PRODUCTION
The abundance pattern of opal in deep sea sediments is closely related to the pattern of productivity of diatoms and radiolarians in the overlying waters.
As shown in figure 2-2, high opal
contents are found around the perimeter of the Antarctic continent, along the equatorial belt in the Pacific, off the coast of
Africa in the Atlantic, and in the northernmost Pacific.
In allof these regions surface wa ters are all characterized by high nutrient content.
These high concentrations are maintained by intense upwelling.
One might ask why it is opal producing plantsthat dwell in these areas of intense upwelling and not, for example, calcite producing plants.
The answer to this question, likemos t of those regarding ecologic interac tions, is not known.
It
can only be said that based on observa tion, if sufficient dissol ved sil ica te is present in surface wa ter, it will be dia toms
that flourish.
Only when silica has been depleted from the water
do other plant types come to the fore.
In this connec tion we canpoint back to figure 1-18, which shows the distribution of N03 andH 4S10 4 in high-la ti tud e surface wa ters.
It is clear from these

47

~

00

Figure 2-2.
The distribution of opal in marine sediments.
Prepared by Ross Heathof Oregon State University.
It must be kept in mind that) since this map shows the
proportion of opal in the sediment) the patterns depend on the rate of accumulationof the non-opal components of the sedime~t as well as that for opal.

distributions that silicate is depleted from the water before nitrate (i.e., high concentrations of nitrate extend much further
equatorward in surface water than do those of silicate).
This
difference must be the chemical imprint of high diatom producti vi ty!

OPAL SOLUTION ON THE SEA FLOOR
The ocean is everywhere undersaturated with respect to themineral phase opal.
Thus once formed the hard parts of diatoms
are immediately subject to solution. As can be seen from the sediment trap data in table 1-3, there is no marked decrease in theamount of opal caught as a function of depth.
Apparently theopaline material suffers relatively little solution as it falls
through the water column.
Rather, most of the opal solution
occurs on the sea floor.
In the paragraphs which follow we will
see that this sea floor solution enhances the contrast between theopal content of sediments forming beneath areas of high and low
opal production.
The extent to which the opal raining to the sea floor dissolves depends on three major factors.
First, it is influenced by
the thermodynamic driving force which is a function of the temperature, pressure and H 4 Si0 4 content of the bottom water.
The
greater this difference the greater the tendency for the opal todissol ve.
Second, it is influenced by the rate of rain of nonopaline material.
This rna terial dilutes the opal in bottom sediments and thereby cuts down the rate at which it dissolves. For agiven thermodynamic driving force the rate of opal solution will
be proportional to the fraction of opal in the sediment. Finally,
it is influenced by the rain rate of opal itself.
The higher therain rate the greater the probabili ty that a grain of opal will
become buried before succumbing to solution.
An analogy will help to clarify this situation.
A botanist
receives a grant to study the process of "leaf detachment."He
hires a research assistant and takes him off to a nearby experimental forest where he has marked off several study plots.
Theassistant is instructed to begin each day by raking one of theplots clean and then to spend the day picking up each oak leafthat falls to the ground within the plot.
These leaves are to be
put into a box for weighing at the end of each day.
On the firstday the assistant chooses a plot containing only oaks. The air iscalm and he has little trouble picking up all the oak leaves whichfall. He decides that he has landed an easy job. The next day heagain picks a plot containing only oak trees.
All goes well forthe first half hour and then the wind picks up, bringing down agreat deluge of leaves.
Try as he may, he can not keep up withthis onslaught, and by the end of the day he finds that he isscurrying about on a carpet of oak leaves.
On the third morning
the somewhat disillusioned assistant comes up with a clever idea.
He chooses a patch within which oaks make up only a small fractionof the trees.
Less work, he conjectures.
Despi te more moderate
winds than on the previous day, he immediately runs into trouble.
Not being an expert on leaf identification, he finds himself pondering each leaf so long that he cannot get to all the leaves
which fall. By noon he notes that oak leaves, buried under leaves49

of other kinds, have missed his immediate attention.
At the day's
end he trudges back to the laboratory carrying his sparse catch
and wondering what new surprises his next day on the job would
bring.
The assistant's first days were analogous to the situation in
the Antarctic Ocean, where it is largely diatoms that rain to thesea floor.
The factor which determined whether or not oak leaves
accumulated was the rate at which these objects fell compared tothe maximum rate at which he could retrieve them.
His third day
was analogous to the more usual situation in the deep sea, where
far more non-opal than opal fragments fall.
In this case his
retrieval rate decreased greatly due to the competition by non-oak
leaves for his attention. Furthermore, the more oak leaves he didretrieve, the more difficult it became to find the few thatremained.
In order to quantify this situation we must define 6 quantities:1)
the rain rate of opal, Ropal
2)
the accumulation rate of opal, Ao~al3) the accumulation rate of non-opallne material, Aother4) the solution rate of opal, Sopal.
5) the solution rate of a hypothetical sediment consistingof pure opal, S *opal
6) the fraction opal in the sediment, f.
These 6 variables are interrelated by 3 equations.
First,
the actual rate of opal solution is related to the hypotheticalrate at which a sediment consisting of pure opal would dissolve by
the fraction of opal in the sediment.
This relationship can bestated as follows:
Sopal=fnS*opal
2-1
The value of the exponent, n, depends on the nature of the solution process. Were only grains sitting on the top of the sedimentcolumn dissolving, then n would be one.
However, as we shall see
if solution occurs mainly in the pores of the sediment then theappropriate value of n is one half.
Second, as sediments accumulate on any given region of the sea floor at nearly uniform rates
over long periods of time, there must be a balance between therain rate of biogenic silica, the accumulation rate of biogenic
silica, and the solution rate of biogenic silica.
This balancecan be written as follows:
Ropal

=

2-2

Aopal + Sopal

The third relationship stems from the definition of f,f

Aopal
Aopal + Aother

i.e.:

2-3

The parameters f and Sooal can be eliminated by simul taneous solution of these three equations yielding the following expression for the ratio of the rate at which opal accumulates inthe sediment to the rate at which biogenic silica rains onto the50sea floor (i.e., the fraction of opal preserved).

J

Aopal

S* 0 pa 1 [Ao paln
1- Ropal
Aopal + Aother

Ropal

2-4

Al though the solu tion to this equa tion is algebraically cumbersome, its main message can be understood by considering two limiting cases.
Fo r sediments rich in biogenic silica we can simplify theequation using the following approximation:
Aopal

= 1

2-5

Aopal + AotherThus regardless of the choice of n:
Aopal

Ropal - S*opal

Ropal

Ropal

2-6

or
Aopal

2-7

Ropal - S*opal=

Such sediments can form only if Ropal is greater than S*opal.
Although we don't have a firm knowled~e of the value of S*opal,
it must be on the order of 5 g/cm 2 lO
years (i.e., much higher
than ordinary opal rain rates).
If so, then in order to achieve
an opal accumulation of 20 g/cm 2 10 3 years, as is observed in someparts of the Antarc tic Ocean, rain ra tes of 25 g/cm 2 1000 years
would be required.
As shown in figure 2-2 most deep sea sediments are very low
in biogenic silica (less than a few percent by weight). For thesesediments,
Ao pal << Ao ther
Wi th this approximation the equation becomes:
Aopal
Ropal

1 _ S* 0 pa I
Ropal

[A o pal J n
Aother

If the exponent n is taken to beequation can be rewritten as follows:
Aopal
Ropal

2- 8

equal

1

=

1 +

to uni ty,

then the2-9

S*opalAotherFor deep-sea sediments wi th accumula tion51

ra tesof less

than 0.52g/cm 103years, S* is much greater than Aother.

Aopal '" Aother
Ropal
S*opalHence:
2-10

For example, for a sediment wi th an accumulation rate of 0.5
2g/cm 10 3 yr and an S* of 5 g/cm 2 10 3 yrs, approximately ten percentof the silica reaching the sea floor would be preserved.
Thus we see that our intuition is confirmed.
Sediments underareas of high opal production will lose only a small fraction of
their opal to solution.
By contrast sediments under areas of low
opal production will experience large fractional losses.
The solution process enhances the contrast between the opal content of
sediments forming under areas of high and low opal productionl
Now let us consider what is the appropriate value for the exponent, n.
Measurements of the dissolved silica content of porewaters in sediments provide us with important clues.
As seen in
figure 2-3, these concentrations increase with depth in the sediment column and gradually approach a cons tant value. *This observation suggests two things.
First, some of the solution mustoccur mainly within the sediment column (the products of the solution process are carried from the sediment by molecular diffusion
through the pore spaces).
Second, the rate limiting step for solution must be the time required for the pore waters to reach thesteady state concentration.
If so, then at any given depth in thesediment the folloWing balance must be maintained:Ds

a2 Cporeaz 2-k (Csat - Cpore )2-11

where Ds is the whole sediment diffusion coefficient**, Csat
and Cpore are the concentrations of silica at "saturation" with
opal and in the pore water at depth z in the sediment, k is thefraction of the residual pore water inequilibrium (i.e., the difference between Csat and Cpore )removed each second.
The
solution to this differential equation is as follows:(C pore - Cbot) = (Csat - Cbot) [l-exp (-z/k/D s )]2-12

where Cbot is the concentration of silica in the bottom water.
This mathematical model predicts an exponential shape for the porewater silica content with an asymptotic approach to the steadys tat e va 1u e .
Th e 0 b s e r ve d prof i 1 e s ( see figure 2- 3) ha ve th is*This steady state pore water value is lower than the saturation
value measured in the laboratory (see figure 2-4).
It must reflecta kinetic balance between the solution of opal and the depositionof some crystalline phase of Si02 or the recrystallization of someclay minera~.
**D s = ~D/ e where D is the molecular diffusion coefficient, <p isthe porosity and e is the tortuosity factor.

52

H45i04 (pmol/kg)

H 45104 (p.mol/kg)

600

o
1/2 EQUIL(4 em)

,

E

E
u

I

-

I-

I-

u

J:

a..

0...
W
(J1

39°N 72 0 W(2.6km)Wooc,...,)

liON 140 0 W(4.9km)o
ATLANTIC

PACIFIC

40"

,

,

,

,

,

"

80'

,

,

,

,

Figure 2-3.
On the left are shown pore water profiles for two cores from theequatorial Pacific Ocean.
The core from lIoN contains 18 percent opal and thatfrom ION contains 13 percent opal. The results were obtained by steve Emerson andDavid Heggie at the University of Washington, using conventional methods forshipboard pore water extraction. On the right is shown a pore water profile for ahemipelagic sediment from the continental margin of the North Atlantic.
The
resul ts were obtained by Fred Sayles of Woods Hole Oceanographic Insti tution,using an in situ pore water extraction device (529a).

~owarz CRYS

,

I

eRYS

UU

en

w 10

-J

HIGH'

AM'OR
SILICAU U..

BiaGE Nrc SIL ICAFROM DEEP SEA
SEDIMENTS

a.

-

~

<l

en 5

-

~

1.L

o
o•
z

I

I

I

I

200

400

600

I

800

1000

1200

H4Si04 (jLmoJ/kg)
Figure 2-4.
Measurements of the solubili ty (in seawater at 3°C)of acid-cleaned biogenic silica separated from deep sea cores collected in central equatorial Pacific.
The measurements were madeby Hurd and his colleagues at the University of Hawaii (522).
Also shown are the solubilities of the minerals quartz, low crystobalite and high crystobalite and the solubility of amorphous silica glass. .

54

basic shape.
The e folding depth in the sediment z* (i.e. thedepth at which the opal saturation and the bottom water concentration difference decreases by a factor of lie) is given by:z* = IDs/k = IDsT

2-13

where T, the resaturation time
waters, is the reciprocal of k.
Solving for T we obtain:

for

the

silicate

content

of

pore

2-14

T

F rom

the profiles for silica-rich sediments shown in figure 2-3,
the value of z* can be estimated to be about 4 cm.
The wholesediment diffusion coefficient for silica in sea water (at 2°C) is
about 2 x 10- 6 cm 2 /sec.
Hence:= 8 x 10 6 sec
or about 100 days.
This implies that if we could by magic replace
the pore water in these sediments with silicate-free sea water it
would take about 100 days for enough opal to dissolve to bring thesilicate concentration lie of the way to the saturation value.
The time required for the pore waters to resaturate with silica would depend on the fraction of opal contained by the sediment.
The time required in a sediment containing 50% opal would logicallybe twice that for a core of pure opal.
As the sediments on which
the pore water studies were performed averaged· only 10% opal, thepore water resaturation time for a sediment of pure opal should be
only about 10 days.
If we call the saturation time for a core withpure opal T * then:
T

*

2-15

T

f

Our equation then becomes:(C pore - Cbot) = (Csat - Cbot) [l-exp (-z/f/DsT*)]

2-16

The flux of dissolved silica through the sediment water interface is given by the relationship:
Flux

=

Ds

~)
dZ

s

~*
f-

z = 0

T

2-17

f1/2

2-18

This yields:
Sopal
S*opal55

The upshot of all this is then that in the case of opal solution
the best choice for n is 1/2.
Let us apply this rule to our botan~
ical assistant.
If he could pick 64 leaves a minute when only oakleaves were present on the ground then his harvest would fall to 45
leaves per minute when half the leaves were from oak and to 32
leaves per minute when one quarter of the leaves were oak.
Solutions of equation 2-4 (using n = 1/2) are shown in figure
2-5, where the fraction of opal which accumulates (i.e., Aopall
Ropal) is plotted against opal rain rate (i.e., RoQal) for several values of the pure opal solution rate (i.e., S*opal).
For
any given value of S*opal the fraction of opal preserved increases with increasing opal rain rate.
What are the appropriate rates for the ocean? An estimate of
the mean ocean opal rain rate can be obtained by dividing the total
amount of silicate dissolved in the ocean (about 18 x 10 16 moles)
by the area of the ocean (3.6 x 10 18 cm 2 ) and by the mean ventilation time of the deep sea (1000 years). *This yields an av~rag~
opal rain rate of 5 x 10- 2 moles/cm 2 10 3 yrs or about 3 g/cm 10years.
This result can be compared with the average solution rate ofopal as calculated from the pore water gradients of H4Si04.
As
shown above, these profiles can be approximated by the equation:(C pore - Cbot) = (Csat - Cbot) [1 - exp (-z/z*)J

2-19

The flux through the sediment water interface is given by:

Ds

~)
dZ

Z

= 0

Z

2-20

*

If we take Ds to be 2 x 10- 6 cm 2 /sec, z* to be 4 cm
and Csat- Cbot to be 400 x 10- 6 moles/kg (4 x 10- 7 moles/ em ~) the flux
becomes 2 x 10- 13 moles/cm 2 sec or 0.4 g per cm 2 10 3 yrs.
Althoughthis estimate of the rate of silicate regeneration is almost afactor of ten lower than the rate of production, two factors must,however, be considered when making this comparison:
1) The cores for which the pore water measurements are shown
come from areas of low opal production.
2) Some of the opal dissolves as it falls through the watercolumn; some also dissol yes on the sea floor before it is stirred
into the sediments by burrowing organisms.
Taking the average rain rate of opal to be 3 g/cm 2 10 3 yrs andthe average" fraction of opal preserved in the sediment to b~ 0.05
we can see from figure 2-5 that the appropriate value of S opal
*As we shall see in chapter 5 the distribution of radiocarbon suggest that the waters of the deep sea are flushed with water fromthe warm surface ocean (transported to the polar regions andcooled) about once each 1000 years.
This in turn implies that all
the water in the deep sea passes through the warm surface ocean
layer about once each 1000 years.
As it does so its dissolved silicate will be converted to opal and rained to the deep sea.

56

-I

<l

Q..

o

0::

.......

-J
<l
0-

.4

o

«

.2

Figure 2-5. The relationship between the fraction of opal fallingto the sea floor which accumulates in the sediment (i.e., Aopall
Ropal) and. the rain ra te of opal (Ropal) for various values ofthe solution rate for a deposit of pure opal (i.e., values of S*).
The accumulation ra te of non-opaline components (Aot~r) is taken to be 0.5 g/cm 2 10 3 yrs.
The units for S* are g1cm 10 3 yrs.

57

would lie between 5 and 10 glcm 210 3 years.
In making this calculation we must keep in mind that because of the wide range of opal
rain rates in the sea the use of averages is not entirely appropriaie. However, it does allow us to get an idea of the magnitudeof S opal.
Thus ,we see that the accumulation of opal varies greatly fromplace to place on the sea floor for two major reasons. First, thep:oduction of opaline silica is concentrated in areas of upwelIlng. Second, the solution process enhances this primary contrast.

DISTRIBUTION OF CALC~TE IN MARINE SEDIMENTS
The distribution of calcite in sediments differs greatly fromthat of opal (see figure 2-6).
While the locations of opal richsediments correspond closely to areas of upwelling, calci te richsediments in the open ocean are found on ridge crests and othertopographic highs.
As the mountains of the continents are drapedwith snow, the mountains on the ocean floor are draped with sediment rich in calcite!
One factor contributing to this difference
in distribution is the geographical uniformi ty of the productionof calci te by organisms in surface waters.
The abundance of theorganisms producing calcite is limited by the availability of nitrate and phosphate rather than by the availabili ty of silicate.
Unlike opal which falls into the deep sea before solution the organic tissue containing phosphate and nitrate is largely oxidized
within the surface sea permitting these two constituents to travelmuch further from their point of upwelling than does silicate.
Itis the combination of the dominance of diatoms in upwelling areas
and the recycling of the ingredients of soft tissue which gi ves
rise to the great difference in the geographic patterns of opal
and calcite production.
Unlike the situation for opal, not all parts of the sea floor
are bathed in water undersaturated with respect to calcite. Whilethe deepest waters in the ocean are corrosi ve toward calci te thewaters which bathe the crests of the mid-ocean ridges are supersaturated with respect to calcite.
The nature of the sediments inthese areas reflect this difference.
Sediments on ridge crests
contain abundant calci te which shows Ii ttle or no solution effects.
By contrast sediments in the deepest parts of the ocean
have lost almost all their calcite by solution. Thus while opal issubject to solution attack everywhere on the sea floor, calcite is
attacked in some places and not in others.
Thus, .unlike the distribution of opal in sediments which reflects mainly the great contrast in opal production from place toplace at the sea surface, the distribution of calcite reflects
contrasts in the extent of solution with depth in the sea.
The depth dependence of the calcite content of sediments fromvarious regions of the ocean is shown in figure 2-7.
As can beseen the boundary between sediments wi th well preserved calci teand sediments free of calci te is not sharp.
A zone of grada tionin the extent of calcite dissolution is found separating calcite
ooze and red clay. We shall see that this transition zone owes its
origin to the finite time required for pore waters to reach saturation with respect to the mineral calcite.

58

Figure 2-6. Map showing the distribution of calcite-rich sedimentson the sea floor.
Sediments containing more than 75% CaC0 3 areshown by the cross-hatched area. This map is based on an Atlantic
map published by Biscaye and his colleagues (489), a Pacific map
published by Berger and his colleagues (491), and an Indian Ocean
map published by Kolla and his colleagues (490).

Table 2-2.
Saturation carbonate ion content of sea water as afunction of temperature and pressure for the minerals calcite(precipitated by coccolithophorida and foraminifera) and for themineral aragonite (precipitated by pteropods).
Saturation C%rbonate Ion Content,
10- moles/kg
Temperature,

24
2
2
2

°c Pressure, atm* Calcite
1
1
250
500

Aragonite

46
48

66
69

71
106

101
152

*A pressure of 100 atm is achieved at close to 1000 meters depth
in the sea.

59

3

.,..

WESTERN
EQUATORIALPACIFIC

I.....
· ••

•

r.

I

0

••

•

•

••

•

-, -

- ,.,
-~4
..
-.If fA•• ••-/,..-

5-·
I ••

.-.

.

j:..

••

EQUATORIAL
INDIAN
looN-IOoS',,CENTRAL
EQUATORIALPACIFIC
5°S-15°N

o

E

~4

:r:
.....
a.
w
o

,• ..
~

••

SOUTHERNINDIAN

~.

••••

•

..e,

Li
..4, • ••~4t••• • :r:-- _.-. . .

... .
•••
•
. .. . -•

..

t~·~

5F-

•

lOa

~.~
u

20

.....
a.

W

I.

o

Ii

40 60 80 lOa
CALC! T E(%)

a

...

,

••••

•

.-

• •
•
• •

-

•

•

-•
• • • -.-.1.
.-- •
.1

• •

,.

•

..
. ·.. ·.

-'~

•

.-

• •••

20

,e

I

\

20 0 -30 0 S...•••.
•• • •..

• - I •• •

:i ·
6··'
f.-

,

EASTERNNORTH
ATLANTIC
15°-35°N

5

,

,

••
• • ••

.....
•..
•
••
•••
• • •
-,......
..: "..
.
.
....
.
,.... "-..
....
.
. •••, ...,, ...•, ,Ilt.....
..,..

3
CJ")

WESTERNNORTH
ATLANTIC
15° - 35°N

:r:
.....
a.
w
o

.. -

••• •

•

3

••
••

•

..
-

•

:r:

I

......•..
•

looN - laos

.....
a.
w

I

-

-

WESTERN
SOUTH
ATLANTIC
28°-34°S

40 60 80 100
CALCITE(%)

a

•

•

.,

•
•

WESTERN
SOUTH
ATLANTIC
15°-17°S

100

Figure 2-7.
Plots of percent CaCa 3 for core top samples as a function of water
depth for various regions of the ocean.
Based on a data compilation provided by
Pierre Biscaye of the Lamont-Doherty Geological Observatory.

Before we continue, it must be pointed out that not just themineral calci te (the stable form of CaCa 3 at any temperature andpressure found within the ocean) but also the mineral aragonite (ametastable form of CaC0 3 under oceanic conditions) is precipitatedby marine organisms.
Since all surface ocean water is supersaturated w~th respect to both calcite and aragonite, organisms generate whlchever form of CaC0 3 they find easier to precipitate. Foraminifera and coccolithophorida choose to form calcite; pteropods
choose to form aragonite.
Although these two mineral phases show
the same basic pattern of abundance in sediments, the boundary between aragonite-bearing and aragonite-free sediments occurs at
much shallower depths than that for calcite.

DEGREE OF CALCITE SATURATION
To understand the relationship between the distribution of
GaCO 3 in sediments and the distribution of biologically utilized
consti tuents wi thin the sea, we need to define the designations:saturation, undersaturation and supersaturation.
For any mineralplaced in contact with sea water, the product of the concentrations of the ions formed during its solution will reach a limit atwhich the solution has reached equilibrium with the given mineral.
This limi t depends on the temperature and pressure (the depth)of the sea water.
In the solution of CaG03, calcium ions andcarbonate ions are formed.
The degree of saturation D is writtenas:D =([Ca++J[C0 3 =J)sea water([Ca++J[C03=J)sea water saturated with calcite

2-21

where the brackets indicate the concentration of the enclosed ion.
Obviously, working wi th measurements on sea water samples
alone will give us only the values in the numerator of this expression for D.
It will not tell us the degree of saturation of
GaG0 3 for the sample.
To determine the value of the denominator,
we must add crystals of calcite to a sea water solution maintainedin the laboratory and wait long enough for the water and the solid
to achieve equilibrium.
The ratio of the oceanic ionic product tothe ionic product obtained from these laboratory experiments gives

D.
For studies in the open ocean, the equation can be simplified.
The calcium content of sea water is nearly constant, itvaries with respect to the salt matrix only by about 0.5 percent
and as a result of salinity differences by only another percent ortyl"lo. *Since the varia tions in CO 3 = ion are by comparis on far
larger, we can assume that the Ca++ ion content is the same
everywhere in the sea and in the sea water samples used in thelabora tory.
Thus our expression for D can be simplified to:*The evaporation-precipitation cycle results in salinity differences from place to place in the sea. The salt is left behind whenwater evaporates, raising the salinity of the residual sea water.

61

D=

2-22
[ C0 3=Jsaturated sea water

D is then simply the ratio of the carbonate ion content in the sea
water sample to the carbonate ion content in the sea water solution saturated in the laboratory at the same temperature and pressure.
The saturation carbonate ion content
varies with temperature and pressure and of course is different for the mineral calcite than for the mineral aragonite.
Aragonite is more soluble
than calcite; thus the equilibrium carbonate ion content of aragonite is always higher at any given temperature and pressure thanthat of calcite (see table 2-2).
The temperature range in the ocean extends from about -loC in
deep and polar waters to about 30°C in tropical surface waters;
the pressure range extends from 1 atmosphere (atm) at the surface
to 550 atm at a depth of 5500 meters (the maximum bottom depth formuch of the sea).
Now CaC0 3 is an unusual sal t:
the colder thewater, the more soluble it becomes!
The difference in the saturation concentration of C03= ions in a surface water sample fromthe warm surface ocean, where the temperature averages 24°C, and asurface water sample from the polar regions, where the averagetemperature reaches 2°C, is about 5 percent (see table 2-2).
Thesolubility of CaC03 also increases with pressure, and pressure increases with depth.
All the water in the deep ocean is cold.
Hence there is no point in considering what happens at a highpressure and a temperature of 24°C;
we only need to know what
happens close to 2°C. Table 2-2 shows that both calcite and aragonite are slightly more than twice as soluble at 500 atm (2°C) as
they are at 1 atm (2°C).
In the deep sea the effect of pressure
is clearly greater than that of temperature.
Moreover, since thetemperature below 1000 meters is nearly constant, variations inthe solubility of CaC03 within the deep water mass are entirely
due to the increase of pressure with depth.
Although much work has been done to pin down the exact solubilities of the mineral phases, calcite and aragonite as shown intables 2-3 and 2-4 and in figure 2-8 uncertainties still remain.

VARIATIONS IN THE CARBONATE ION CONTENT OF SEA WATER
One important ques tion remains.
How does the carbonate ion
concentration vary from place to place within the sea?
With thisinformation in hand, we could see how the degree of supersaturation of sea water with respect to the CaC03 minerals calcite andaragonite varies from place to place within the sea.
To understand these variations, we must probe more deeply into some of thechemical reactions taking place in the ocean.
The C03= ion content of sea water must vary with the total
amount of dissolved inorganic carbon.
We have seen in chapter 1that the total dissol ved inorganic carbon content is about 20%
higher in deep North Pacific water than in warm surface water.
As
we shall see, the range in carbonate ion content is far large rthan this (a factor of 6) and in the opposite sense (high in warm
surface water and low in deep North Pacific water).
The source of these variations is changes in the fraction

62

Table 2-3 Estimates of the solubilities of the minerals calcite and aragonite in surface sea water (35.0 0 /00 salinity).
Investigators

Institution(year)Ref.
CALCITE

Ingle andCoworkers*

Oregon state( 1973)(459)

Broecker and
Takahashi**Lamont-Doherty(1978)

Plath andCoworkers*
Morse andCoworkers*
Sayles+Adopted

2

4.5
4.7

44
46

(463)

2

4.8

47

Oregon state( 1980)

(466)

25
5

4.7
4.8

46
47

Miami
(1980)

(465)

25

4.4

43

Woods Hole(1980)

(464)

2

4.4

43

2

4.8

47

for this book

25

ARAGONITE

86

Yale( 1976)(461 )

Plath andCoworkers*

Oregon state

(467)

25

Morse andCoworkers*

Miami( 1980)

(465)

25

6.6

64

25
2

6.6
6.8

66

2

6.9

67

Berner*

25

8.8

11.4

5

III

85

(1980)

1.43 times calcite sol.++
Adopted for this book

* Based on laboratory equillbratlons.
** Based on lysocline depths and in situ saturometry.
+ Based on pore water measurements.
++ Solubility ratio based on thermodynamic arguments.

63

64

Table 2-4. Estimates of the pressure dependence of the solubility of calcite.

Investigators

Institution(year)

Ref"

MethodPytkowicz andFowlerOregon State(1967)( 457a)Laboratory
equilibration-31

Hawley and
Pytkowicz

Oregon state(1969)( 458)

Laboratory
equilibration

-33

Ingle

Oregon State( 1975)( 46 0)Laboratory
equilibration

-42

Yale
( 1969)(458a)

Partial molar
volumes-39

DuedalOregon State(1972)( 458b)Partial molar
volumes-43

Broecker and
Takahashi

Lamont-Doherty (463)( 1978)

Millero andBernerSaylesWoods Hole(1980)
Adopted for this book

~V

*

--

cm 3/mole

Lysocline depths -37
and in si tu
sa turometry(464)

Pore watermeasurements

-42
-37

_ z
_ 0
V/Rrr
*[C0 3 -J sa t
[C0 3 -J sa t e-~
where AV is the difference between the volume occupied byCa++ and C0 3 = in solution in sea water and in the latticeof the mineral calcite.
Table 2-5. Charge balance in sea water:

the excess cation charge
is balanced by the dissociation of carbonic acid (H 2 C0 3 ) into bicarbonate (HC03-) and carbonate (C0 3 =) ions.

Cation
Na+
K+
Mg++
Ca++
E

PositiveMassmoles/kg
0.470
0.010
0.053
0.010

-

Chargeeq/kg
o. 470
0.010
0.106
0.020
0.606

Anion
ClS04=BrE

HC0 3+
C0 3 =
E'

64

Nega tive
Massmoles/kg
0.547
0.028
0.001

Charge
eq/kg
0.547
0.056
0.001

-

0.604

-

0.002

-

0.606

30

70

110

150

0 ....- .....,....--....- -.....- - - - . - - -.....---.

~

2

E

PORE WATER
MEASUREMENTS
BY SAYLES

~

""""

.c
.....

3

a..
Q)

0

4

ADOPTED IN
THIS BOOK

5

6 .....

--.a.---~--.....- _....._.........._.. . . . .

Figure 2-8. Plots of the saturation C0 3 = ion concentration versus
depth in the sea based on the analysis by Broecker and Takahashi(463) of lysocline depth and saturometer data and measurements by
Sayles (464) of the C0 3 = ion content in pore waters extracted insitu from deep sea sediments.

65

of the total dissol ved carbon that is in the form of carbona teion.
Dissol ved carbon is present in the ocean in three forms:
dissolved carbon dioxide gas (C0 2 ), bicarbonate ion ( HC0 3-) andcarbona te ion (CO 3 =) •Carbon is one of the few elements in the~cean.that can exist in ions havi~g different charges:
as HC03lon Wl th a charge of -1;
as C03 - ion wi th a charge of -2.
Theamount of gaseous C02 in the ocean is very small and will for thesake of simplicity be neglected in the discussions which follow.
The ra tio of CO 3= to HCO 3- plays an importa"nt role in the
sea's balance of electrical charge.
The major contributors of
positively charged ions (cations) in the sea are Na+
K+ Mg++and Ca++ (see table 2-5).
To find the total number 'of u~its orpositive charge per kilogram of seawater, we mUltiply the moles/kg
of each cation by its charge and total the resul ts.
This yields
O. 470 equi valents of charge from Na+, 0.010 equivalents from K+,
0.106 equivalents from Mg++, and 0.020 equivalents from Ca++.
The
total is 0.606 equivalents of positive charge per kilogram. *Since sea water can have no overall charge, the number of equivalents of negatively charged ions (anions) must add up to exactly
the same value.
If we take the three major anions (chloride(CI-), sulfate (S04=), and bromide (Br-»), we find that theircharges add up to 0.604 equivalents of negative charge per kilogram, so we are short 0.002 equivalents of negative charge perkilogram of sea water.
This difference between the sum of thenegati ve charges is balanced mainly by the dis sol ved carbon ions
HCO 3- and CO 3='
The ratio of these two anions to one anothervaries in such a way that charge is balanced.
When more negative
charge is needed to balance the cations present, HCO 3- is converted to C03=;
when less is needed, C03= is converted to
HC03-'
The amount of excess positive charge balanced by bicarbonate and carbonate ions is a small part of the total but, as weshall see, a very important residual that holds the key to thecarbonate ion distribution within the sea.
In the discussion which follows we neglect the element boron.
The presence of boron in sea water requires a small correction in
all the calculations outlined here, for it also exists in two
forms (B (OH) 3 and B (OH) 4 - ) . In this firs t look at carbonate chemistry we keep things more simple by considering a hypotheticalboron free ocean.
To understand what determines the ratio of carbonate ions tobicarbonate ions in a given unit of sea water, we must recognize
two restrictions.
First, the total amount of dissolved carbon
present in. the water cannot be changed as a result of our manipulation of the charges. There is a fixed number of carbon atoms andthese must exist in either bicarbonate or carbonate form.
Second,the sum of all the bicarbonate and carbonate charges must justbalance the excess cations so that the sea water maintains its
electrical neutrali ty.
Respecti vely, these two restrictions canbe stated mathematically as:*Chemists measure alkalinity in units of equivalents/kg. An equivalent is equal numerically to a mole.
One mole of Na+ ionscarries one equivalent of positive charge; one mole of 804= ionscarries two equivalents of negative charge.

66

2-23
where LC02 stands for total dissolved inorganic carbon, and2-24
where Alk is the alkalini ty of the water (that is, the exces s
pos i ti ve charge to be balanced by CO 3 == and HCO 3- ions).
The
alkalinity is given by:
[AlkJ
If

==

[Na+J + [K+J + 2[Mg++J + 2[Ca++J + ..•-[CI-J - 2[S04=J - [Br-J-···

we

subtract

equation

2-23

from

equation

2-25
2-24,

we

find

that:

2-26
Thus the carbonate ion content of any unit of sea water sample isequal to its alkalinity (excess positive charge) minus its total
dissolved carbon content.
SUbstituting this result in either ofthe original equations, we obtain:
2-27
Here we see that the bicarbonate ion content equals twice thetotal dissolved carbon content minus the alkalinity.
So to understand how the carbonate ion content of sea water
varies from place to place, all we need to know is how the alkalinity of the water and how the total dissolved carbon content of
the water vary.
As we have already seen, the total dissol ved carbon conten tof sea water varies because plants extract carbon from surface sea
water and because the remains of these plants (or of the animals
they support) sink into deep parts of the oceans where they arelargely destroyed.
Two processes are at work:organic tissue
formation and destruction and hard-part formation and destruction.
The formation of organic tissue consumes carbon and hencechanges the LC02 of the water.
This carbon loss has no effect on
the alkalini ty of the water because none of the ionic concentrations appearing in the equation defining alkalinity change.
However the incorporation and release of ni trogen by organic matterdoes make' small changes in alkalini ty.
When organically boundnitrogen is released during respiration, the N03- ion so produced
adds to the anionic charge and reduces the alkalinity of the deep
water.
Correspondingly, the removal of the NO 3- ion to form
organic matter increases the alkalinity of surface water.
Removal of CaC03 from the water changes both concentration of
the total dissolved carbon and the alkalinity. Organisms use carbon to form CaC03 and thus remove it from the water. The alkalinity of the water changes because when CaC0 3 is formed, Ca++ ion(one of the contributors to the net positive charge) is removed.
Each mole of CaC03 formed results in the extraction of 2 moles of
positive charge; thus the creation of CaC03 changes the alkalinity67

of sea water by just twice the amount that it changes the carbon
content. The formation of the third type of particulate material
opal, affects neither the alkalinity nor the total carbon content
of the water.
In chapter 1, we estimated that for every four C atoms thatare removed from surface water as organic tissue, one C atom isremoved as CaC0 3 .
Table 2-6 shows the total effect on sea waterof adding or sUbtracting organic debris plus CaCD 3.
From thetable, we see that for every five C atoms used (four in organic
material and one in CaC0 3 ) one Ca atom 1s used.
This uses up two
positive charges and correspondingly reduces the alkalinity.
Foreach 4 moles of carbon incorporated by plants about 0.6 moles ofnitrogen are incorporated.
This raises the alkalini ty by 0.6
equivalents partly compensating for the reduction of alkalinity
due to C.aC.o 3 forma tion.
For the combined organic-C.aC03 debris,the change in alkalinity is equal to 28% the change in total dissolved carbon.
If the composition of the debris being formed anddestroyed were constant everywhere, then the carbonate ion content
of sea water would follow some very simple rules .
Although thisis not quite the case, it provides a good first approximation of
the operation of the system.
To see how the CO 3 = ion content varies, let us start wi thsurface water that is phosphorus-free.
Remember that this watercontains the minimum amount of carbon.
Such phosphorus-free sea
water has an alkalinity of 2.20 x 10- 3 equiv/kg and a total dissolved carbon content of 2.00 x 10- 3 moles/kg.
The carbonate ioncontent in this water must be equal to 2.20 x 10- 3 equiv/kg minus
2.00 x 10- 3 moles/kg, or 0.20 x 10- 3 moles/kg.
Now let us assume that our sample of surface water is transferred to the deep sea. While in residence there, it receives theoxidation products of enough organic tissue and the solution products of enough CaC0 3 to raise its total dis sol ved carbon 3by 0.20
x 10- 3 • So instead of having a LC02 con~ent of 2.00 x 10- moles/
kg, we have a LC0 2 content of 2.20 x 10- moles/kg.
If the organic tissue and the CaC03 are destroyed in roughly the same propositions in which they are formed (so that for every increase of 5moles in LC02, there is an increase of 1 mole in Ca++ and an increase of 0.6 equiv in N03-), then the alkalinity increase is 0.06
x 10- 3 moles/kg.*
Since surface water has an alkalinity of 2.20x*Actually the ratio of CaC03 dissolved to organic tissue respired increases with depth.
This can be seen in figure 2-9 where alkalinity(normal ized to a salinity of 3S. 0 0 /00) is plot ted against EcO 2 concentration (normalized to a salinity of 35.0 0/00).
On this diagram
points for temperate and Antarctic surface water and for various deep
water types are shown.
The expected s lopes are calculated from thefollowing equation:lS
6C orgMlk
tlCCaC03--= 2
lOS
tlEco 2tlC CaCO 3 + 8Corg--

The second term in this equation is the correction for nitrate. As canbe seen the deep waters of the Indian and Pacific oceans fall along atrend consistent with a ratio of one mole of CaC03 dissolved per one
mole of organic residues respired. On the other hand, the line connect
ing warm surface water and mean deep water has a slope equivalent to onmole of CaC03 dissolved per four moles of organic residues respired.

68

Table 2-6. The impact of particulate matter formation on thechemical composition of sea water.
Component Formed

Constituent
ChangeOrganic
TissueCalciteAggregate

-4.0
0.0
-0.6
+0.6

-1.0
-1.0
0.0
-2.0

-5.0
-1.0
-0.6
-1.4

6[ECO*]
6[Ca+ ]
6[ N03-]
6[A]

0')

c.o

Table 2-7. Idealized carbonate chemistry of average surface and deep water.

Water
Type

Carbonate
Bicarbona te
GaseousIon (HC0 3-),Ion CC03=)
CO2
10-3 mo l es /kg 10- 3moles/kg 10-3moles/kg

Total
Dissolved
Inorganic
Carbon (C0 2+
HC03+C03=),10- 3moles/kg

Alkalinity(HC0 3 - + 2C03=)
10- 3 eq/kg

Surface

0.01

1.80

0.20

2.01

2.20

Deep

0.03

2.14

0.06

2.23

2.26

-----r-----.....- ---.. . .

~----.._---NORT I NOlANAND PAC' FIC
BOTTOM WATER

2500 .....

.....-__..

~I MOLE CoC0 3
IMOLE ORG C

2450

t7'

SURFACE~~oX

V

I MOLE CoC0 3:L

4 MOLES ORGCz

6ALK =lx2_1x~= 29
6IC02 55 105 •

~

c:{

~ ,

WATER

.......

..J

2 105 .

ANTARCTIC

2400

tT

_1 ~= 93

6ALK =1 2
6!C02 2

~

~

.,

~

2350

~ NORTH ATLANT IC@DEEP WATER (2.5°C)

DEEP
WATER

SURFACEWATER

x WSBW

ANTARCTIC
S. INDIAN... S. PACIFIC+ CPOW~ NIBW
I:i NPBWo NADW@•• N.PACIFIC• S.ATLANTIC• N.ATLANTIC

2250L------J~----1._----L------L---------

1900

2000

2100

2200

IC02

N

2300

2400

(p.mol/kg)

Figure 2-9.. Plot of alkalini ty agains t ~CO 2 concentra tion (b othnormalized to a salinity of 35.00%0) for major surface anddeep water types in the world ocean. The solid lines show thetrends expected for CaC03 to organic C input ratios of 1:4 and1: 1.
Based on resul ts obtained as part of the GEOSECS program(406 ).

70

10- 3 equiv/kg, our deep water must have an alkalinity of 2.26 x
10- 3 equi v/kg.
The difference between alkalini ty and total dissolved carbon in our deep water sample is 0.06 x 10- 3 moles/kg.
Despi te the fact that deep water has 10 percent more total dissolved carbon, it has only about one third the carbonate ion content of surface water. This is so because more CD2 is released todeep water by the oxidation of organic tissue than C03= ion by
the solution of CaC 0 3 •The excess carbon dioxide combines wi th
carbonate ion to form bicarbonate (C02 + C03=+ H20 ~ 2HC03-)'
The HC0 3 - ion content of the deep water is 2.14 x 10- 3 moles/kg,
which is, of course, higher than that in surface water. Table 2-7
swnmarizes these calculations.
In this table the dissolved C02(but not the borate contribution) is included.
The vertical distribution of carbonate ion content (calculated wi thout any of the approximations used above) along northsouth sections in the world ocean (calculated with all species included) is shown in figures 2-10 (the western central Atlantic
Ocean), 2-11 (the western central Indian Ocean), and 2-12 (theeastern and western Pacific Ocean). A quick look at these figures
shows that the large difference in carbonate ion concentration obtained in the simplified calculation is also found in the realocean.
Between 30 0 N and 30°3 the concentration in surface water
generally is above 200 x 10- 6 moles/kg, reaching nearly as high as
300 x 10- 6 moles/kg.
By contrast, the values in deep water range
from 65 to 115 x 10- 6 moles/kg.
A closer look at the deep water
values reveals that they decrease from the Atlantic Ocean to theIndian and Pacific Oceans (a distribution consistent with that ofdissolved oxygen and of nitrate).
Further, the basal part of themain thermocline in all three oceans shows a carbonate ion content
minimum.
This minimum corresponds to the minimum in dissolved
oxygen content (again consistent with expectation).

SPAC.IAL VARIATIONS IN THE CaC03 SATURATION OF SEA WATER
Having established the distribution of carbonate ion in theocean we can now compare it with the saturation carbonate ion content for calcite and for aragonite.
To illustrate this relationship, the carbonate ion concentration versus water depth curve fora station in the South Atlantic is compared in figure 2-13 withthe saturation carbonate ion versus depth curves for calcite andaragonite. As can be seen, the carbonate ion saturation curve foraragonite 'crosses over the observed carbonate ion concentration
curve at a depth of 3.4 kilometers.
The calcite saturation curve
crosses over the in situ curve at a depth of 4.3 kilometers.
Thewaters above these cross-over depths are supersaturated with respect to this mineral phase. The waters beneath the cross-overdepth are undersaturated with respect to calcite.
The difference
between the two values(i.e.,[C.03=Jin situ
[CD3=Jsat)isa measure of the tendency for the calci te to dissol ve .
Positivevalues indicate supersaturation and negative values undersatura tion; the greater the difference, the greater the extent ofsupersaturation or undersaturation.
To show how the depth of the saturation horizon for calcite
varies within a given ocean and between oceans, vertical sections

71

E2
~

I

l-

n..

w

o 3

WESTERN
ATLA~TIC

110

C03

100

100

/

9~IOOJ-LmOI/kg~
100/ \""-- 90-"~45'--'"""'----a..-.a..-----.&-.a..---a...---a.-....

30
S

20

10
0
10
LATITUDE

20N

5 ---.----.......-----.......-----20
20

S

10
0
10
LATITUDE

N

Figure 2-10.
The left hand diagram shows the distribution of carbonate ion concentration with depth along a north-south section in
the western basin of the Atlantic Ocean.
The minima in carbonate
ion content water at a depth of about 0.75 kilometers to the northand to the south of the equator are the result of higher productivity supported by upwelling in the equatorial zone.
The wedge oflow carbonate concentration water at the bottom of the South Atlantic is due to the invasion of Antarctic Bottom Water. Based on
the resul ts obtained as part of the GEOSECS program (406).
The
tracks along which this and the sections which follow were drawn
can be found on the foldout map.
Figure 2-11.
The right hand diagram shows the distribution ofcarbonate ion concentration with depth along a north-south sectionin the western Indian Ocean.
Because this ocean has no northernsource of deep water, respiration effects are large and thecarbonate ion contents correspondingly low.
Based on results
obtained as part of the GEOSECS program (406).

72

I

f0...

w
03

I

70

f-

EASTERN

0..

PACIFIC

0

w

CO=3

. ,umo I/kg

80

80

WESTERN

44PAC IFIC
CO=3,umol/kg

5

10
S

0

10

20

LATITUDE

5
10
S

30
N

0

10

20

LATITUDE

30

40N

Figure 2-12. Sections of carbonate ion concentration with depth
along north-south sections in the eastern (to the left) and western (to the right) Pacific Ocean. Again, because there is no deep
water source in the North Pacific, the effects of respiration become stronger going from south to north, leading to a northward
decrease in carbonate ion content in the deep waters.
Based onresults obtained as part of the GEOSECS program (406).

73[CO;] (jLrnOl/kg)
0

0

40

80

120

0

0

200

160
0

240

280

0

<:e

0
0

0

0
0

0
0

0

~

0
0

E

ATLANTIC

0

-

0

a.

0

0

0

0~18°S-31°W

~

J:
fo-

0

w

0
0
0,
0

0

ARAGONITE
SATURATION

0
0

0
0
0

0

0
5
Figure 2-13.
Plot of carbonate ion concentration versus water
depth for a station in the western South Atlantic.
Also shown are
the saturation carbonate ion concentrations increase wi th depth
mainly because of the pressure effect (calcium and carbonate ionsoccupy less volume when dissolved in sea water than when locked in
CaC.O 3).
The concentration of carbonate ion remains fairly constant through the North Atlantic Deep Water mass.
It is higher inthis mass than in the underlying wedge of Antarctic 'Bottom Water
and in the overlying tongue of Antarctic Intermediate Water.
These results were obtained as part of the GEOSECS program (406).

74

running north to south through the three major oceans for the property ~C03 (i.e. ,[C0 3 =J1D ~ [C0 3=J sa t)
have been prepared (see figures 2-14, 2-15 and 2-16). In these diagrams thezero contour represents the saturation horizon.
The contours fornegative values represent undersaturated water, and those forpositive values supersaturated water.
The saturation horizon isdeepest in the western Atlantic Ocean (- 4500 meters), intermediate in the western Indian Ocean (- 3500 meters) and shallowest in
the North Pacific « 3000 meters).
Figure 2-17 shows a similar section constructed for aragonite
in the western Atlantic. * Consistent with its metastability (andhence higher solubility), the depth of the saturation horizon foraragoni te in the wes tern Atlantic is abou t 1000 meters shallower
than that for calcite.
In the Pacific and Indian Oceans the satura tion horizon for aragoni te is loca ted wi thin the main thermocline.
The depth of the saturation horizon determined from sedimentary evidence is broadly consis tent wi th that determined from carbonate ion measurements.
As shown in figures 2-14, 2-15 and 2-16,
the depth at which the effec ts of solu tion are first seen in thesediments (termed the lysocline) is about the same as the depth or
the saturation horizon.
As both depth estimates are sUbject touncertainties of several hundred meters, the differences between
the saturation horizon depth and the lysocline depth seen in these
diagrams lies within our ability to measure them.
Below the lysocline, the percent calcite in sediments accumulating on the flanks of the mid-ocean ridges drops from the valueof 90 percent characterizing sediments above the lysocline to justa few percent 300 to 1000 meters below the lysocline.
Figure 2-18
shows the trend of the calci te content of sediments as a func tionof bottom depth for several regions of the sea floor.
These diagrams also show the water depth interval over which the difference[C03=Jinsitu[C0 3=J sa tchangesfroma to 10 x 10- 6
moles/Kg. ~can be seen from these diagrams that an undersaturation of 10 x 10- 6 moles/kg is enough to dissolve almost all the
calcite falling to the sea floor.
In order to provide yet another independent es tima te of thedepths at which the minerals calcite and aragonite begin to dissol ve, an oceanographer named Mel Peterson came up wi th a clever
new approach (473).
He hung calcite spheres at various depths on
a mooring set by physical oceanographers interested in measuring
the speed and direction of currents in the deep sea.
By carefully
weighing these spheres before and after deployment in the sea, he
was able to gain a sense of this variation of the rate of calcite
solution with depth in the sea.
His resul ts showed considerable
sca tter, bu t yielded the trend shown in the left panel of figu re
2-19.
Since Peterson's demonstration of the utility of this method, other inves tiga tors have done similar experiments.
The mos telaborate of these was carried out by Honjo (475).
They deployed*As the saturation horizons for aragonite in the central Indian andPacific Ocean are shallower than 600 meters,presenting corresponding diagrams.

75

there is no point in

-

--

E

~

10

WESTERN INDIAN~C03 CALCITE
P. moll kg

1----0_0
CALCITE. / LYSOCLINE

4 •••••••••••----201----

WESTERN ATLANTIC

10-------~CO~ CALCITE
p.mol/kg

6~-Ao--""'---.l""-"-"""'-_.a-.---._-""

s

30

20

10

0

10

20

N

LATITUDE

5'---.......-------&.--------'
10
0
10
20
LATITUDE
Ns

Figure 2-14. The left hand diagram shows a section in the western
central Atlantic of the vertical distribution of the difference
between the in situ carbonate ion concentration and the saturation
carbonate io~concentration for the mineral phase calcite.
Where
this difference is zero, the waters are saturated with calcite.
Where it is negative, calcite dissolves.
The rate of solution is
proportional to the magnitude of this difference. The rise in thesaturation horizon to the south is related to the invasion of bottom water from the Antarctic carrying lower carbonate ion content
than does the deep water formed at the northern end of the Atlantic.
The depth of the lysocline for calci te as determined fromsediment studies is also shown.
These resul ts were obtained aspart of the GEOSECS program (406).

Figure 2-15. The right hand diagram shows a section in the western central Indian Ocean of the vertical distribution of the difference between in situ
- carbonate ion concentration and the saturation carbonate ion concentration for the mineral phase calcite.
The depth of the lysocline for calcite as determined by studies ofthe deep sea sediments is shown for comparison. Based on data obtained as part of the GEOSECS program (406) and on the solubilityof calcite and recommended by Broecker and Takahashi (463).

76

0

WESTERN
PACIFIC~C03 CALCITE

0-

EASTERN
PACIFICE-

fLmol/kg

I-

0

-E

~C03 CALCIT

.oX:

I

a..

fLmol/kg

-

.oX:

J: 3

0

t-

a..

w

----

w

0

0

~ CALCITE

LYSOCLINE

4

CALCITE-10 LYSOCL INE-10

5 ...........- - - & . - - . 1 - - - & . . . - - - - 1 . - - - 1

s

10

o

10

20

LATITUDE

5 L-_J.==::::I:::::::::..I-_...L_--L....J
10
o
10
20
30
40

30

N

S

LATITUDE

N

Figure 2-16. Section in the eastern (on the left) and western (onthe right) Pacific showing the vertical distribution of the difference between the in situ carbonate ion concentration and thesaturation carbonate ion concentration for the mineral phase calcite. It CM be seen that the saturation horizon rises to depths
less than 1000 meters in the north Pacific.
The depth of the lysocline as determined from sediment studies is shown for comparison.
Based on data obtained as part of the GEOSECS program (406)and on the solubility of calcite recommended by Broecker and Takahashi (463).

77

20

WESTERN
ATLANT IC

t::. C0:3 ARAGON IT E

2

fL mol/kg.....-.,.

Ex.
.....,

10

J:..... 3a..

ARAGONITE~LYSOCLINE

W
0

0
-10

4

-20

-30:
-40=
5 __--""__--..__

s

30

20

-

-

.

.

~

10

1oo_

1oo_

~

...

-

0

LATITUDE

N

Figure 2-17. Section in the western central Atlantic showing thevertical dis tribu tion of the difference between the in si tu carbonate ion concentration and the saturation carbonate-ron concentration for the mineral phase aragonite.
The saturation horizon
is at a shallower depth than that for calcite (see figure 2-14).
The depth of the lysocline for the aragonitic shells of pteropods
based on sediment studies is shown for comparison.
Based on dataobtained as part of the GEOSECS program (406) and on the solubility of aragonite recommended by Broecker and Takahashi (463).

78

CALC ITE IN SEDIMENT (%)

a

20

40

60

80

100
E

:r:

.-

3.5

0W

..lll::RIO GRANDE RISE
30°5
WESTERN ATLANT ICOONTONG-JAVA PLATEAU
100N-100S
WESTERN PAC' F rC

3.0

i=
0W

o
:E

4.0:E

m

t5. 0 ~

o

.-o
6 4.5

o

t-

- . . 0 _ . . & . - _........---.;.---11-_..&.-

E

~

3.5

.-:r:o..

E
3.0 ~EAST PAC' F IC RISE
17°S
EASTERN PACIFIC

I

t-

a..

4.0 w

W
0 4.0
~

.-.-0 4.5
0
m

0

2

~~C03=. 5.0

10 p.mol/kg

:=0
0

co

0
40
20
60
80
100 0
20
40
60
80
100
CALCITE IN SEDIMENT (%) CALCITE IN SEDIMENT (%)

Figure 2-18. Generalized plots of the calcite content of material
from sediment core tops versus water depth (on the flanks of midocean ridges and plateaus).
While indi vidual data points show afair amount of scatter~ enough measurements have been made to permit the mean trends shown here to be defined.
The depth at which
the calci te content of the sediment begins to decline is des ignated as the lysocline.
Below this depth, solution becomes progressively more intense until nearly all the calcite which rainsto the sea ,floor dissolves.
This progressive intensification ofsolution reflects the increase in the degree of undersaturation
with water depth beneath the saturation horizon.
The arrows show
the water depth interval over which the difference between the insi tu and saturation carbonate ion concentrations changes by 10 x
10- 0 moles/kg.
The top of each arrow is at the lysocline.
Thedata on which these generalized diagrams are based are shown infigure 2-7.

79

small pumps at three depths on a mooring and pulled water overvarious types of calcite and aragonite.
The pumps eliminated theeffect of differing current flows.
Each pump pushed water acrossthe crystals at the same rate.
Their results are also shown in
the right panel of figure 2-19.
While more detailed measurementsof this type are needed~ the results of those carried out to dateare certainly consistent with the distribution of calcite andaragonite in sea floor sediments.
The resul ts of a more recent deployment of Honjo' s pumps in
the Panama Basin are summarized in figure 2-20 (476).
In this experiment both calcite and foraminifera shells were placed in thechambers.
After exposure the calcite crystals were reweighed andthe foraminifera shells viewed for breakage.
As can be seen in
the figure the chambers at water depths less than 3 kilometers
showed some weight loss (3 to 8%) and some breakage of foraminifera shells (presumably from solution damage).
The chambers located at 3.8 kilometers showed much greater calcite weight loss
(24 to 33%) and far more foraminifera shell damage (15 to 23% of
the shells were broken).
These changes are nicely reflected by
the calcite content of sediment samples taken at various depths inthis area.
Sediments from depths less than 3 kilometers are highin calcite (>80%).
Below 3 kilometers a rapid drop in calci tecontent is seen.
The deepest sample (at 3.5 kID) had only 43% calcite. As can be seen from figure 2-16 in this region of the eastern Pacific, the waters between 2 and 3 kilometers depth are justslightly supersaturated (the Honjo data suggest that perhaps they
are just slightly undersaturated) and that at about 3.4 kilometers
they become undersaturated.
The data in figure 2-16 also suggestthat at depths in excess of 3.4 kilometers the degree of undersaturation increases wi th depth (Honjo' s data suggest that this increase commences at about 3.0 kilometers depth).
Thus the two
sets of measurements agree to within a l!.C03= value of about 5 x
10- 6 moles.
One other line of evidence must be discussed.
Almost two
decades ago Peter Weyl invented a technique for the direct measurement of the undersaturation and supersaturation of calcite (oraragonite) in sea water (469).
His procedure was as follows.
Hefilled a tube wi th crystals of calcite (or aragoni te) and buried
the tip of the glass electrode of a pH meter in these crystals.
He then pumped sea water through the crystals at a high enoughrate so that the pH in the water leaving the tube was exactly thesame as that in the water entering the tube.
He then stopped theflow of tbe water through the tube and in so doing isolated abatch of sea water within the crystal pores. He then observed thechange in pH of this water with time (if the water was initially
undersaturated~ the pH rose and if the water was initially supersaturated, the pH fell).
Examples of the responses found for undersaturated sea water are shown in figure 2-21. As shown, the pH
approaches a constant value asymptotically.
Assuming this value
to be that for saturation with respect to the mineral phase present in the column, Weyl showed that the ratio of the ini tial to
the final CO 3 = ion concentra tion could be obtained from the pH
results using the following relationships:

80

,......--.,,.......-.......- -.......- -... a

0 ....- - . . . - - - - . - - -....._---.

-E 2-

ATLANTIC
OCEAN(32°N-55°W)

PACIFIC
OCEAN(19°N-156°W)

2x ARAGONITE

~

.::I:

0-

0

\

ARAG.

ex

~LYSOCLINE

4-+- LYSO.

3

w

E

oX

• CALC ITE~ """FOR CALCITE~LYSO.'x

5

CALC.
CALCITE SPHERES

5""" -fTEROPODS

x

cocco. FOR AM. REAG.

6

0.1
0.2
0.3
0.4
0.0
2
SOLUTION RATE (mg /cm yr)

6
50
75
100
WEIGHT LOSS IN 79 DAYS (%)

a

25

Figure 2-19. Results of mineral exposure experiments performed on
deep sea moorings (473).
The experiment depicted on the left was
conducted by Peterson of the Scripps Institute of Oceanography in1965 for a period of 250 days using calcite spheres hung directlyin the sea.
The experiment depicted on the right was conductedby Honjo and his coworkers at the Woods Hole Oceanographic Institution in "1976 for a period of 79 days (475).
In this experiment, the rna terials (coccoli ths, foram shells -' reagent calci te,and pteropod shells) were held in chambers through which water
was drawn by a pump.
The depth of the lysocline for the respective mineral phases as established through studies of the sediments are shown for comparison.

81

CALCITE WEIGHT LOSS (%)
AND
BROKEN FORAM I N I FERAL TESTS (%)

00

5

10

15

20

25

30

PAC I Fie 0 CEAN
PANAMA BASINo SURF. SED. CoC0 3 CONT.

-E-2• CALC. WEIGHT LOSS•BROKEN FORAM TESTS..:It:

:I:
ICL
lJJ
Q

3

4 ....._ ......- -....._ _......_ ..........._ ......._ _.........
100
90
80
70
60
50
40
SURFACE SEDIMENT CaC03 CONTENT (%)

Figure 2-20. Comparison of weight loss (solid circles) and percentage of broken foraminiferal tests (solid triangles) for an insitu mineral and shell exposure conducted in the Panama Basin byHonjo and his coworkers at the Woods Hole Oceanographic Insti tution (476) with the percentage of calcium carbonate (open circles)
for surface sediments taken at various depths in the same area.

82

60

, CALCITE

\

•••

-.

45

••I.

........
C

E

~

w 30

,• •

-l-~

.,

15

...•
•

•

• •
••

z
0
lL.

LL.
LL.

LL.

0

z

0

LL.
LL.

• -.

•
••

• .. .. .

~

••

LL.
LL.

\.

0

0

••

•• ST 341•.., 4.5 km
-5

0

0

• •

0

LL.
LL.

-10

z

z

...

0

LL.

0

• •

•
•

0

LL.
LL.

.

•

LL.

0

0

,

••

•
•
ST 3384.2 km

•• ••
-5

-10

~EMF(mV)~EMF(mV)

Figure 2-21. Electromotive force (i.e., potential) difference between the pH electrodes immersed in calcite (on the left) and inaragonite (on the right) against time using the Weyl-Ben Yaakov insi tu saturometer.
The sharp offsets occur when the pump flushes
fresh seawater through the column.
When the pump goes off, the pH
drifts toward the saturation value.
These resul ts were obtained
during the GEOSECS program.

83

[co 3=JFinal[H+JInitial[ C03=JInitial[H+JFinal

2-29

Further, if the CO 3 = ion concentra tion in the input wa ter were
known, then the difference between the saturation and initial
C03= ion concentrations could be obtained:

H: -1][ CO 3 =] sat - [ C0 3= ] I = [ [
Jr
[H JF

[C 0 3= ] I

2-30

Some years after Weyl invented this clever technique, Sam Ben
Yaakov adapted it to work within the sea (471).
He built a unit
which could be lowered from a ship into the deep sea.
It had apump which went on and off in a fixed time cycle and it had a taperecorder which kept track of the electrode potential difference
and hence of the pH.
The records shown in figure 2-21 were obtained by this unit.
These measurements of the depth of the saturation horizon arecompared in figure 2-22 wi th those obtained by assuming that the
sediment lysocline represents the saturation horizon.
The agreement is quite good.
One might ask why an ocean-going saturometershould be used to measure the solUbility of calcite and aragonitewhen such measurements can be made in the laboratory.
The answer
is that the laboratory measurements prove to be very difficul t.
It is not easy to maintain sea water at low temperatures and highpressures without changing its alkalinity due to interactions withthe container or its LC0 2 content due to biological activity or to
C02 release from the container.
Experiments in the deep sea circumvent the need to maintain all these properties.
FACTORS CONTROLLING THE RATE OF CALCITE SOLUTION

Can the same arguments made for the rate of opal solution beapplied to calcite?
While the answer to this question is yes,
there are several major differences between the solution of thesetwo phases.
The first has to do with the driving force for solution.
Since the saturation silicate content of sea water (withrespect to opal) is so much larger than the ambient silicate content of deep sea water, the driving force for silica dissolution
is nearly constant over the entire deep sea floor (i. e., S* isnearly constant).
We have seen that for calcite the driving force
varies wiqely, it ranges from zero at the lysocline to as high as
30 x 10- 6 moles/kg difference in CO 3= ion concentration in the
deepest parts of the ocean.
Thus the variabili ty in S* for calcite becomes a major factor in determining the distribution of
calcite in sediments.
Another difference is that the pore water resaturation time
for calcite appears to be about 4000 times smaller than that foropal!
We know this from the response seen using the Weyl-Ben
Yaakov "saturometer" (see figure 2-21).
Rather than taking 10days (-8 x 10 5 sec) as calculated from the silicate versus depth
profiles in sediment pore waters, only a few minutes (-2 x 10 284Co3' (,umol/kg)..--..

E 2

~

-

--...

J:
l-

n.
w
0 3

x

4

x

••

\
x

x

~•

•

5

Figure 2-22 Plot of saturation CO 3 = ion concentrations obtainedusing the Weyl-Ben Yaakov in situ saturometer (crosses) and by assuming that the lysocline represents the·saturation horizon (solid
circles) •The curve drawn through these points has the exponential shape demanded by thermodynamic theory.
Z in the equation
indica tes wa ter depth in kilometers.
This summary was made by
Broecker and Takahashi (463).

85sec) are required for the pH to adjust in a calcite-filled column.
As shown in equation 2-13, the relationship between the characteristic resaturation depth and the characteristic resaturation
time is:z* = IDs T

2-31

Hence,
z*calcite

Tcalcitez*opal

Topal

~

0.016

2-32

The resaturation depth for pure opal is about 1 cm. Thus that forpure calcite must be on the order of 0.16 millimeters (160 microns) .
This fast response time allows calcite to dissolve as fast as
opal despite the fact that the difference between the steady porewater concentration and the bottom water concentration for opal(~300 x 10- 6 moles/kg) is 30 times larger than that for calcite(at the base of the transition zone between calcite rich and calcite free sediment the value of ~CD3 is about 10 x 10- 6 moles/kg).
As shown in equation 2-17, the flux is given by:
Flux = (C-sat - Cbot)

J~s

2-33

Hence the 30 fold difference in concentration gradient is compensated by the 4000 fold difference in resaturation time!
There are two other factors which might influence the rate ofsolution of calcite. First, fluid dynamicists tell us that at the
sediment water interface there is a barrier to molecular transport. We can approximate the resistance imposed by this boundary
layer by thinking of it as a "8 tagnant film" through which ionspass only by molecular diffusion. The resistance of the film will
be proportional to its thickness.
Although it is not possible todetermine the effecti ve thickness of this film, the best guesses
place it in the range 30 to 3000 mi crons .
The point is that, if
the thickness lies at the low end of this range, the resistancewhich it poses is small compared with that posed by the diffusion
path within the sediment (averages about 100 microns in length).
On the other hand, were it to lie at the high end of the range,
then its influence would be greater than that posed by the diffus ion pa th in the sediment.
These si tua tions are shown diagramatically in figure 2-23.
Since silicate has a characteristic diffusion path of several em in the sediment, its solution rate would
not be significantly altered by even a film of 3000 microns (i.e.,
0.3 cm) in thickness.
Whether it is the "stagnant film" just above the sedimentwater interface or the diffusive pathway within the pore waters of
the sediment which constitutes the limiting resistance for calcite
solution has implications for the distribution of calcite on thesea floor.
Were the stagnant film the limiting resistance, then86

COi (jLmol/kg)

3

C0 (,umol/kg)

-

en

c
0

~

u

E

~

w
u
~

LL

a::

W

I-

Z

~

0

a::

LL

W
U

Z
<1: -400

I(f)

-

0

-800

Figure2-23. Hypothetical plots of C03= ion content versus distance away from the sediment water interface.
In the upper two
diagrams the boundary layer is assumed to exert a resistance toionic transport equivalent to that of a stagnant film 15 microns
in thickness, and in the lower two diagrams it is assmned to exert
a resistance equivalept to that of a stagnant film 300 microns inthickness.
The left hand diagrams are for sediments wi th 100%
calcite an"d the right hand diagrams are for sediments with 25%
calcite.
If the pore water resaturation time is inversely proportional to the calcite content of the sediment, then the resaturation depth will be proportional to the square root of the fractionof calcite in the sediment.
Thus a two times greater resaturation
time would be found for the 25% calcite sediment than for the 100%
calc i te sediment (i. e., an e-folding depth of 600 microns rather
than 300 microns).
In each diagram the saturation C03= ion
content is shown.

87

the rate of solution would depend on the factors controlling the
thickness of this film (i. e., current veloci ty and turbulence inthe abyssal waters).
The rate of solution would not depend on the
calci te content of the sediment.
On the other hand, if the diffusive pathway wi thin the sediment pore waters constitutes thelimiting resistance then the solution rate would depend on the
calcite content of the sediment but not on the water flow rate andturbulence in the abyssal waters.
Thus it is important that welearn more about the magni tUde of the resistance to moleculartransport posed by the dynamic boundary layer just above the seasediment interface.
During the period this book was being written Santschi andhis coworkers at the Lamont-Doherty Geological Observatory devised
and tested a method to measure the resistance posed by so-called
benthic boundary layer (696).
Using Alvin, a research submersible, Santschi deployed and recovered flat plates of alabasterfrom the sea floor off Central America.
Based on the weight loss
during the period of four days residence on the sea floor he wasable to show that the stagnant boundary film thickness over thealabaster was about 120 microns.
As the bottom currents in thearea of deployment were considerably higher than average this result is likely smaller than the ambient deep sea value.
Once thismethod has received further testing and wide deployment it may bepossible to determine whether it poses a significant resistance tocalcite solution.
As proposed by Emerson of the University of Washington and
Bender of the University of Rhode Island there is another way inwhich the solution process for c~lcite differs from that for opal.
If organic material raining to the sea floor is stirred into thesediment column before being eaten by scavengers or bacteria, thenthe C02 generated from its destruction will be added to the porewaters. Before this C02 can diffuse out it will be neutralized by
the solution of calcite.
The products of this reaction would thendiffuse through the pore spaces out into the overlying sea.
This
process would lead to calcite solution at all depths in the sea,even above the lysocline (519).
The ratio of organic residues to calci te entering the sediment column is cri tical for this process.
If there were one ormore moles of organic carbon per mole of CaC03 then it would havethe potential to dissolve all the calcite.
On the other hand, if
the ratio is less than 0.1, then the destruction of organic matter
would be secondary in importance.
As we pointed out in chapter 1,
sediment traps deployed near the sea floor generally collect particulates with an organic residue carbon to CaC03 carbon ratio
close to uni ty.
How much of this organic debris gets mixed into
the sediment before being eaten? As organic matter accumulates indeep sea sediments at only 1% the rate at which it rains toward
the sea floor, virtually all is eaten.
The fragments of organic
matter falling to the sea floor provide the sole source of foodfor bottom-dwelling organisms, thus one might expect these organisms to be adept at reaching and consuming newly fallen morsels
before they are buried.
It would therefore be surprising if asubstantial fraction of the organic matter were inadvertently
stirred into the sediment by the very organisms seeking it for88

food!
On the other hand, one might argue that the organisms whichburrow through t~e sediment do so in search of food which they
have purposely mlxed down as a way of preventing it frombeing
eaten by organisms which cannot burrow, or that some fraction of
the organic material is inedible by bottom-dwelling animals and is
eventually consumed by bacteria living within the sediment. Right
now we simply do not know what fraction of the raining organic
debris gets buried before it is eaten.
Another question that arises for calcite is whether its
dissolution rate can be assumed to be proportional to the AC03=
value.
Indeed, evidence exists that this may not be the case.
Measurements carried out by Keir in Berner's laboratory at Yalewith calcite grains suspended in a rapidly stirred solution haveshown that the solution rate is proportional to the fourth powerof AC03 (479).
If these results apply to t:Q.e sea floor, then therate of solution is 16 times higher for a AC0 3 of -20 x 10- 6
moles/kg than for a AC03 of -10 x 10- 6 moles/kg!

THICKNESS AND SHAPE OF THE SUBLYSOCLINE TRANSITION ZONE
If we put aside for the moment the possible resistance
imposed by the boundary film at the sediment-water interface, thepossibility that the relationship between the solution rate ofpure calcite and the AC03= value is other than first order andthe influence of the organic matter stirred into the sediment,
then we can easily ascertain the factors controlling the width of
the zone separating sediments experiencing little calcite solution
from those for which nearly all the calcite has dissolved and theshape of the calci te content versus wa ter depth curve acros s thezone.
If the calcite content of the sediments within this zone isassumed to be at steady state (i.e., conditions in the ocean haveremained cons tant for many thousands of years), then the calci tesolution rate, S, can be equated to the difference between therain rate, R, and accumulation rate, A. Hence:
S

= R -

2-34

A

This equation can be rewritten:

S R [1 - LfL=

2-35

where fL is the fraction of calci te in the sediment above thelysocline and f is the fraction of calci te at some depth lying
within the transition zone.
As shown above, if the ra te of solution is limi ted by thetime required for pore water saturation and if this reaction is
first order, then:

2-36

89

or

2-37

s = ~C.03

Hence the shape of the calcite content versus depth curve is givenby the equation:
~CO=
3

ftii

f
_s_

2-38

T*

Mos t places in the sea the value of ~CO 3 = increases
stant rate with depth beneath the lysocline.
Hence:

at

a

con-

2-39
where hL is the lysocline depth and h the depth ofwithin the transition zone.
Our equation then becomes:~hR~

a JD;J:

interest

2-40

1 _ f( I-f r ) ]fL(l-f)[While this equation cannot be sol ved conveniently for f, graphs
can be made showing how f varies with ~h for various values of a,R, and fL.
Ds and T* are assumed to have the same value
everywhere on the sea floor (i. e., T * = 120 sec and Ds = 2 X
10- 6 cm 2 /sec).
The curves in figure 2-24 show how the transitionzone would look for several calcite rain rates for a value of fL= 0.9 and a value for the constant a of 10 x 10- 6 moles/kg km(both are typical for much of the deep sea).
Under these conditions:~h= RO.lf] 0.25 km
J~f [1 - O.9(1-f)

2-41

As can be seen, the entire range of thicknesses for the transition
zones observed in various parts of the ocean (see figure 2-7)
could simply be the result of a four fold range in the rain rateof calcite.
Indeed the transition zone thickness differencesfrom region to region do roughly parallel changes in rain rate!
A similar eq ua tion can be wri tten down for the si tuation
where a stagnant film limits the rate of accumulation.
In this
case:
~h =

R ~z

aD

1 _ f ( 1-f I,) ]2-42

fL(l-f)[Taking the same values as above for fL and for the constant a, a
calcite rain rate of 1 g/cm 2 10 3 yrs~ a value for the diffusivityof C03= in bottom water as 1 x 10- cm 2 /sec and a stagnant film

90

0.0
0.0 I

1.0
0.0

0.8
I

A

1.0

I

w

w

z

-J

Z

0.5

...J
U

u
en

~
~
II

STAGNANT FILM
LIMITATION~~g
RCALCITE =2 cm 2 10 3 yr

~1= 1500microns

I

I

~

0.5

oen

0

.......

I

~

~

<.0

I

r*= 120s~~

I

PORE WATERRESATURATION
LIMITATION

-E

-E

0.2

'CALCITE
0.4
0.6

~

1.0

"

1.0

~

~

<l

<l

R= I g /cm 2 103 yr

2-25
1.5"

,

,

,

,

,

,

I .5

PI

,

,

,

,

Figure 2-24. Plots of calcite content of sediments as a function of depth below thelysocline as calculated from the pore water saturation model.
The diffusion rate of
C0 3 = ion in the sediments is assumed to be 2 x 10- 6 cm 2 /sec, the difference between the saturation and bottom water carbonate is assumed to increase at the rateof 10 x 10- 6 moles/kg per kIn and the pore water saturation time T* is assumed tobe 120 seconds.
Figure 2-25.
Plots of the calcite content of sediments as a function of depth beneath the lysocline for the stagnant film model compared with that for the porewater resaturation model.

,thickness,

~z,

~h

of 1500 microns, the equation becomes:[1 - O.~(i:f)JO.48

km

A comparison between the calcite versus depth curve for this modeland that for the pore water saturation case is given in figure
2-25.
As can be seen, the shapes are very different.
One model
predicts a long tail of low calci te sediments, and the other arather sharp base for the transition zone. The observations (seefigure 2-7), while showing considerable scatter, appear to be morein tune with the shapes given by the porewater resaturation model
than by the stagnant film model.
Before leaving this subject, one last point must be made.
The relationship between the calcite content of a sediment and theextent of calcite solution that it has experienced is not what intui tion might lead us to believe.
The si tuation is depicted infigure 2-26.
In this example, above the lysocline the accumulation rate of calci te and of other components of the sediment arein the 9 to 1 proportion in which they rain onto the sea floor.
This sediment has 90% calcite.
At a dep,th below the lysocline
where ~C03= reaches a value of -3 x 10- 6 moles/kg, 5/9 of thecalci te dissol ves, yet the calci te content of the sediment isstill 80%!
At a depth where ~C03 = reaches a value of -6x10- 6
moles/kg the S* value for calcite is twice as high.
However, as
the calcite content of the sediment will be lower, the actual solution rate will not be twice as high.
As shown in the diagram,the solution rate will rise from 5 to 8.
This will lead to asolution of 89% of the calcite but will yield a sediment of 50%
calcite content !
At a depth where the ~CO 3= reaches -9 x 10- 6
moles/kg the rate of solution cannot rise much above what it was
at a ~C03= of -6 x 10- 6 moles/kg (it can only go from 8 to 9 andthen no calcite will be left to dissolve).
Hence the rise must bebalanced by a drop in the fraction of calcite in the sediment.
Inthis situation the following relationship holds:(~C03=)depth 1(~C03=)depth

2= [fdepth

122J /fdepth 1

2-43

Hence:
2-44
In our example:
fdepth 2 = (6/9)2 0.50 = 0.23
As shown in figure 2-26, a sediment wi th 23% calci te would haveexperienced 97% calcite solution!
So we see that the percent
calci te in the sediment drops far more s lowly wi th the extent ofsolution that one might guess.

92

Ll C0 == + 5 fl: mol
3kg

9LlCO = -3 ~mol
3
kg

9
-

9
90%
CALCITE--LYSOCLINE LlCO~=_6fLmol
kg

LiCO = =-9 ~mol
3
kg 8

9

9

I

I

4
80%
CALCITE

I
0.3 I
23%

I

50%
CALCITECALCITE

Figure 2-26. A diagrammatic view of changes in the driving force
for calcite solution, the rate of solution and the percent calcite
in the sediment w~ th water depth.
In each example the right handarrows gi ve the rain ra te and accumulation rate of non GaGa 3
debris and the left hand arrows the rain rate and accumulation
ra te of calci te.
The wavy arrows represent the solu tion rate of
calcite.
It can be seen that the fraction of calcite in the sediment gives a misleading view of the fraction of the raining
calcite which dissolves.

93

VARIATION OF SEDIMENT TYPE WITH TIME
The interaction between the factors controlling the distribution of opal and calcite in the sediments forming on the sea floor
and the great tec tonic processes opera ting in the earth lead to ara ther interes ting sequence of sediment type wi th depth in thesedimentary column at any given point on the sea floor.
Geophysicists have shown that the ocean crust is moving away from thecrests of mid-ocean ridges at the ra te of a few centimeters ayear.
New crust is constantly being generated to fill the void
crea ted by this movement (see figure 2-27).
As the newly formed
crust moves away from its point of origin on the ridge crest, it
accumulates more and more sediment.
Indeed, the thickest sediments are found furthest away from the mid-ocean ridges.
In addi tion to this la teral motion, there is a vertical motion. *The mid-ocean ridge crests have a more or less constantelevation, projecting to within about 2500 meters of the ocean
surface.
The ocean floor well away from the ridge crests lies
about 6000 meters below the sea surface (actually, by this point,
500 meters of sediment have accumulated and the water depth is only about 5500 meters).
As the ocean floor moves away from theridge crests, it drops about 3000 meters in elevation.
Thus atthe time any particular portion of crust was first formed, it wasbathed in CaC03 supersaturated water.
However, as the crust moves
away from the ridge crest, it eventually drops below the calcite
saturation horizon.
As it does so, CaC03 accumulation will slow
and then cease. The CaC03 already accumulated will then be buried
by a layer of red clay (figure 2-27).
We would expect that if aboring were made through the sediments in any region of the sea
floor which currently lies below the sa tura tion horizon, the redclay encountered at the top of the core would eventually give way
to carbonate-rich sediment.
This sequence has indeed been foundin numerous deep sea drillings.
The movement of the ocean floor will often carry a segment of
the crust across the equator (see figure 2-28).
Since opal deposition today is very rapid along the equator in the Pacific Ocean,
we would expect to find a silica-rich horizon in the sedimentary
record of any crustal segment which made such a crossing.
Only
during the period when the crust lay beneath the high productivity
zone, would it accumulate opal-rich sediment.
The expected pattern of sediments with depth is shown in figure 2-28.
This pattern has been found by during the course of the deep-sea drilling
program.
The composition of the sediments forming at a given point on
the sea floor often show short-term cyclic changes as well as thelong term changes associated wi th sea floor spreading.
These
short-term changes have been shown to be in tune wi th the waxing
and waning of large ice masses which existed in what is now Canada
and Scandinavia during glacial periods.
An example of suchchanges is provided by the calci te content of sediment from theequatorial region of the Pacific Ocean.
As shown in figure 2-29,*The lithosphere near the ridge crests is quite hot.
As it moves away,
it cools; as it cools, it contracts.
This contraction produced a drop
in elevation of the sea floor with age.

94

{}

o

o

Alumino-Silicate Detritus

o

{}

{)

+ CaCOJ + Opalo{)

Su pe rsa tu rated
CaCOJ Preserved

SATURATION HORIZON
Undersaturated..
..

.. . .

Figure 2-27. Sequence of sediment types accumulated by the greatli thospheric pla tes as they move away from the cres t of mid-ocean
ridges.
The first sediment to be depos i ted is rich in iron andmanganese oxides, which are products of hydrothermal circulation.
At a distance of more than a few kilometers from the spreading
center, the' sediment no longer receives hydrothermal material andis dominated by calcite falling from the surface. At a distance by
about 1000 kilometers from the crest the plate subsides below thesaturation horizon for calcite, and calcite no longer accumulates.
Beyond this point, continental detritus dominates the sediment.
Cores bored through sediment capped with red clay encounter buried
calci te ooze and then a thin layer of iron- and manganese-richsediment just before entering the underlying basalt.

95

0~

~

EQUATORIAL ZONE OF HIGH PRODUCTIVITY'~CaC03="-I--

IllFe +Mnl

Figure 2-28. Sequence of sediments accumulating on a lithosphericplate which crosses the high productivi ty equatorial belt in its
movement away from a ridge cresta
The numbers relate the map
locations to the corresponding sedimentary sequences found in borings made at these points.

96

Co C0 3 (%)

o

20

40

Or--.,.....--y--r--........-

E

60

80

..............~r-~ .....

E

-

u

u

I

I

I-

l-

W

W

lL

0..

o

o

800

V24-624834m
03°04'N
153°35'W

Figure 2-29. The content of CaCO 3 versus depth in two cores fromthe central equatorial Pacific.
Both were obtained at depths below the lysocline and have experienced calci te solution.
As wewill show in chapter 9, fluctuations on the depth scale of 100 em
are related to glacial to interglacial climate cycles. The longer
term change in core V24-62 reflects either a gradual deepening of
the lysocline or a thickening in the calcite transition zone beneath the lysocline.
These results were obtained by Hays and his
coworkers at the Lamont-Doherty Geological Observatory (499).

97the GaCO 3 con tent of cores in this region flue tua ted wi th time
These varia tions could, of course, resul t from ei ther changes i~
the relative rain rates of materials making up the sediment coresor from changes in the extent of solution of the calcite.
Studiesof these cores show that the latter 1s the dominant cause for theobserved changes.
These studies involved measurements of the ratio of the fragments of foraminifera tes ts to whole foraminifera
shells and also of the ratio of the shells of benthonic foraminifera (organisms which live on the sea floor) to the shells ofplanktonic foraminifera (organisms which inhab it the upper wa terof the ocean).
The utili ty of the first of these indices is obvious.
As solu tion of the shells proceed s, the chambers whichmake up an individual test tend to become disconnected (just as
they did in Honjo's chambers).
Hence, the higher the ratio offragments to whole foraminifera shells, the greater the extent ofsolu tion.
The secorrl method is useful because the shells of benthonic foraminifera are thicker and more rugged than the shells
produced by their planktonic cousins.
They are thus more resistent to the attack of waters undersaturated with respect to calcite.
The ratio of whole shells of benthonic foraminifera towhole shells of planktonic foraminifera increases with the extent
of dissolution.
An example of the time variation in the latterratio is shown in figure 2-30.
These studies show that the lysocline was deeper in the Pacific Ocean during times of glaciation
than during times of interglacia tion such as we are experiencing
today.
The oceanic lysocline also shows changes on the time scale ofmillions of years.
Evidence for these changes is found in thelong borings mad eat over 500 si tes on the ocean floor.
As mentioned above the transi tion from overlying red clay to underlying
calcite-rich ooze was found in many of these borings. The age ofthis transi tion can be obtained from the fossil assemblage in theuppermost calcite bearing sediment.
On the basis of this age anda knowl edge of the ra tes of sea floor sUbsidence* the eleva tionof the sea floor at the time of transi tion can be determined.
Each such point provides a paleolysocline elevation. The temporal
history of the lysocline for the three major ocean basins obtainedin this way is shown in figure 2-31.

MANGANESE NODULES
In add i tion to s il ica te detri tus, calci te, opal, am a traceof organic. matter, deep sea sediments contain concretions of ironand manganese oxide which the oceanographer calls manganese nodules.
These objec ts Ii tter the sediment surface in abou t 50% of
the areas of the deep sea floor.
Nodules are also frequently encountered at depth in deep sea cores.
A photograph showing an
area of dense nodule coverage is shown in figure 2-32.
A cross*As mentioned above, the oceanic crust drops .in elevation due toshrinkage caused by cooling. As the rate time dependence of heatloss (and hence of cooling and shrinking) can be calculated frombasic principles, the elevation change during any time interval in
the history of a given plate can be calculated.

98

CaC0 3 (%)°_>1....
, _7...5

'°..°

05..

x.::-x'"x,'" '"

'"

X
I

50 x,
>f

?<

x
\

X

X

100

X EQUA.

E-

X, PACIFIC

CJ

:I:

t-

a..
w
o

"

~ CORE 60

150xI

x
I

X
I

,

X

200

>f

~

-J
<{

u<{
--l
C)

0:::

lLJ
l-

250

x_ - -xZ--- -50100
BENTH ICS (%)

Figure 2-30. Plots of CaC03 content (solid circles) and percent
benthic forams in total whole foram assemblage (crosses) as afunction of depth in the sediment for a core from the equatorial
Pacific Ocean.
It can be seen that, the percentage of CaC03 wasmuch lower during the last interglacial period than during thelast glacial period and that the percentage of the shells of benthic forams in the coarse fraction sieved from the sediment wasmuch higher.
This combination is the fingerprint for changes
brought about by the solution effects.
Sediments which have suffered no solution effects have only a few percent benthic shells
in the total foram fraction.
During the last interglacial, one
sample shows up to 75% benthics (i.e., 3 benthic shells to every
one planktonic shell)!
To achieve this proportion, more than 99%of the planktonic shells mus t have been reduced to fragments bysolution.
These data were obtained by Arrhenius and his colleagues on materials obtained during the Swedish deep-sea expedition (496).

992r----r----.,--~-_..--..,.....-_
2

E

E

-

3'::;

t-a..

t--

oX

J:

I

4Q.
W

W

0
0

o
o

U
U

5U

INDIAN
60

u

ATLANTIC.....-.-...&-----a_ _.........._ - - L_ _..L.-_--'6

50

lOa

150

0

MILLIONS OF YEARS AGO

50

100

150

MILLIONS OF YEARS AGO

2 r----r----.,r---~--,.---r----..

=.E 3
I

t-a..
w

o
o

U

U

5

PACIFIC
6.....-.-.....L..-----l"-_........_ - - & ._ _. . L . - _......
o
50
lOa
150

MILLIONS OF YEARS AGO

Figure 2-31. A reconstruction of the depth at which the CaC03
content of sediments reaches 20% (te~med the calcite compensation
depth, or CCD). The reconstruction is accomplished by finding theage (and depth) of the transi tion from CaCa 3-rich to CaCa 3-poor. sediment In a large number of long borings made as part of the
Deep Sea Drilling Program.
The transi tions are assumed to haveoccu~red as the sea floor subsided beneath the CCD.
A correctionis then made for the extent of subsidence between the time of thiscrossing and today.
These reconstructions were by Van Andel ofStanford University (508).

100

section through a nodule is shown in figure 2-33.
Their mode oforigin remains somewhat of a puzzle.
The elements of this puzzle
are as follows:
1) As iron and manganese are both extremely insoluble in sea
water, it is not clear how these oxides become concentrated into
concretions inareas of the sea floor remote from continental
margins and ridge crests.
2) Various methods of radioactive dating give concordant
growth rates for nodules averaging about 3mm per million years.
As manganese nodules have radii of a centimeter or more, they must
be several million years old!
Where the sediments in nodule
fields have been dated, sedimentation rates of several meters per
million years are usually obtained.
How then does a nodule remain
atop the sediment during its long period of growth?
During its
lifetime a typical nodule must have worked its way up through ten
meters or more of sediment.
The first of these difficulties can be overcome when one considers that both iron and manganese are very soluble in their reduced form.
Mos t sediments around the margins of the continent
become sufficiently reduced at depth of several centimeters permitting the mobilization of Mn and Fe.
This mobility allows these
metals to escape back to the sea. Thus even though very insoluble
under the oXidiZing conditions of most sea water, iron and manganese atoms could, in a series of water-to-sediment-to-water jumps,
eventually work their way to the remote deep sea, where a stable
home in a manganese nodule (or in oxidized sediment) can finally
be found.
The second difficulty is not so easily overcome.
As nodules
are more dense than the sediments upon which they rest, gravity
cannot be called upon to keep them "afloat" on the ever-thickening
sediment column.
One favori te hypothesis claims that worms whichlive in the sediment every so often nudge a nodule and lift it
through the ever-encroaching sediment.
As no one has ever observed this process in action, its credibility is marginal. Other explanations such as the action of tidal-wave-induced bottom currents are even less palatable.
The implausibility of the mechanisms proposed to keep the nodules afloat on the sediment-water interface has led some investigators to doubt the validity of the figures for the growth rate of
nodules or those for the rate of sediment accumulation.
Beyondthe supporting arguments based on the concordancy of the various
methods used to determine these rates, a quite impressive argument
can be made in defense of the ra tio of the ra te of nodule growth
to the rate of sediment accumulation.
If the process of nodule
formation has reached steady state (as it must), then for each new
nodule which nucleates, an old one must become buried in the sediment.
The nodules so buried could either accumulate as plums inthe sediment pudding, or be dissolved away.
As the nodules foundburied in sediments show no evidence of "corrosion", it is likelythat the fate of most old nodules is permanent burial.
If so,
then the ratio of nodule growth rate to sedimentation rate can beused to predict the densi ty of nodules (i. e., the plums) in thesediment (i.e., the pudding).
For example, in an area where the101

Figure 2-32. Photograph showing a field of unusually closely
spaced manganese nodules on the bottom of the Antarctic Ocean.
The mean nodule diameter is 6 centimeters. Photograph provided by
Charles Hollister of the Woods Hole Oceanoghraphic Institution.

102

F1gure 2-33. Photograph or a cross-section cut through a manganese

nodule with two growth centers.
The altered volcanic material at
the center served to nucleate the nodule's growth.
Crude growthrings representing temporal changes in texture and composition
can be seen.
This nodule has a radius of 2 centimeters.
At agrowth rate of 3 rom/l0 6 yrs, it must have been born about
7,000,000 years ago.
Photograph provided by Sorem and Fos ter or
Washington State University.

103

sediment accumulation rate is three meters per million years, andwhere the mean age of the nodules on the sediment surface (asmeasured by their radius and growth rate) is three million years
the number of nodules found in the upper 9 meters of the sedimentcolumn should equal the number found on the surface.
In otherwords, the previous generation of nodules must lie buried in three
million years' worth of sediment.
For an analogy to this situation, consider a group of spacepioneers living on a large glacier covering the polar region of a
distant planet.
After realizing the futili ty of shoveling aside
the ever-accumulating snow, these colonialists reconcile themsel ves to lifting their bodies and their belongings out of each
new layer of snow.
They also find that they need not bury theirdead; the next snowfall will do the task for them.
After 1000
years of continuous habitation with a steady population of 10,000,
their resident archeologist, out of shear boredom, digs a deep
trench in an attempt to find relics of their arrival.
Finally,when his trench has reached a depth of 300 meters, he succeeds.
Having made note of the artifacts of his forefathers' arrival, heponders the number of bodies he uncovered during his digging.
He
correctly calculates that, as his trench covers 0.1% of the areaof the colony, 10 people should on an average be living within its
bounds at anyone time.
The ave rage lifetime of these colonis ts
was 100 years.
Thus over the 1000 years of occupation, about 100
people should have died and become buried in the snow on the siteof his excavation.
The first generation of pioneers lay in thelowest 30 meters of his excavation, the second in the overlying 30
meters, and so forth.
The members of the eleventh generation currently dot the surface (working from birth to death to keep theirheads above the ever-accumulating snow).
While marine geologists cannot dig trenches on the sea floor
to see whether the appropriate number of "dead" nodules lie buriedin each meter of sediment, they can punch out many, many deep sea
cores.
By comparing the number of nodules which by chance wereencountered within these cores with the number encountered on thetops of these cores, an estimate of the ratio of "dead" nodules
per cubic meter of sediment to "live" nodules per square meter of
sediment surface can be made.
The resul ts of such a study (221)are consistent with the prediction made from the mean age of surface nodules and the mean accumulation rate of the sediment.
The
number of nodules in a sediment layer several meters thick is
about the same as
the number of nodules on the surface!
Thus, "at this point, the worms appear to have the upper hand.
If each year worm nudges were to lift a nodule 3 microns, then ina million years it would be raised by the required 3 meters!
Unfortunately, we have no worm nudge tracers.
So the puzzle remains
an annoyance to those who study the sea floor.
One last question that might be raised about manganese nodules is how these objects manage to capture the small amounts ofdissolved Fe and Mn which reach the deep sea floor.
It turns outthat the implication implicit in this question, namely, that nodules attract more than their share of manganese and iron, is simply not correct.
As shown in table 2-8, the sediments adjacent to
the nodules accumulate manganese at an even greater rate than the104Table 2-8. Comparison of Mn accumulation rates in nodules

and in sediment. These rates are obtained by multiplying
the accumulation rate of the sediment or growth rate of anodule (normalized to the area of sea floor it covers) by
their respective Mn contents (444).
Ocean

Whole ocean
North Atlantic
South Atlantic
IndianAntarctic
South PacificNorth PacificNodules

Sediment

g/ cm 2 10 6 yrs

g/cm 2 10 6 yrs

0.6(5)*
0.4(1)

1.2 (30)*-

1.5 (8)0.6 (1)1.1 (7)0.4 (4)0.2(1)
0.8(3)

1 .. 4 (2)1.4 (8)*Numbers in parentheses indicate number of nodules orcores upon which the average is based.

105

nodules themselv~s!
Boudreau and Scott (695) have suggested that
the Mn accumula tlng on nodules is supplied from the dis sol ved Mn
i~ bottom water and that the rate limiting step for this accumulatlon is the stagnant boundary film discussed above.
Santschi andhis coworkers (696) show that at the site where their alabasterpIa tes were 6 deployed this may well be the case.
The growth ra teof 20 mm/IO years calculated from the Mn content of bottom waterand the measured film thickness of 100 microns agrees wi th thegrowth rate of the nodules in this area.

SUMMARY
Ocean sediments are dominated by alumino silicate debris (redclay) from the continents and opal and calcite hard parts formedby marine organisms.
The rain ra te of clay debris decreases awayfrom the continental margins.
Opal accumula tes mos t rapidly beneath areas of strong upwelling.
As all ocean water is undersaturated with respect to opal, solution of this mineral occurs everywhere.
By contrast, calcite is produced more uniformly over theocean surface, bu t it falls into water which is in some places
supersa tura ted and in others undersa tura ted.
In any given ocean
basin, calci te is abs ent in the deepes t sediments bu t is present
in great abundance on ridges and plateaus which project above thesea floor.
These fac tors lead to strongly contrasting sedimenttypes.
The deep sea floor generally is lined wi th red clay (>90%
alumino silicate debris).
The flanks of the mid-ocean ridges areblanketed with sediments rich in calcite.
Belts of opal-rich sediment are found in the Antarctic Ocean, along the equator in the
Pacific Ocean and along some continental margins.
The extent of opal and calcite solution on the sea ,floor depends on the degree ofundersaturation of the bottom water, therain ra te of these phases, and the rain rate of the dilu ting material.
In both cases, solution occurs primarily in the porespaces of the sediment from whence the produc ts of solu tion diffuse into the overlying bottom water.
C02 generated by organisms
living within the sediment may contribute to the solution of calcite below the lysocline and even lead to partial solution above
the lysocline.
The benthic boundary at the base of the water column may impede calcite solution but not the rate of opal solution.
The movement of the lithosphere away from the mid-ocean ridge
crests and to greater depths determines the long-term sequence of
sediment types on any given portion of the sea floor.
On a shorter time scale, the composition of ocean sediments is influenced byclima tic changes.
For example, during glacial times the calci tedissolution in the deep equatorial Pacific was less intense at any
given depth than it is today.
Nodules of iron and manganese oxide dot the sediment-waterinterface in many parts of the ocean.
Similar nodules are also
found buried in marine sediments.
These objects nucleate on
sharks teeth, volcanic fragments and other hard fragments which by
chance reach the sea floor and tend to grow for millions of years
before succumbing to the incessant rain of sediment.
The great
mystery associated with these objects is the process that keeps
them afloat so long atop the accumulating sediment blanket.

106

PROBLEMS
2-1

If a water sample from 4000 meters depth in the North Pacificis returned to the laboratory and placed under the following
conditions, in each case write an expression based on thesimplifying assumptions used in the text for the ratio of the
new carbonate ion content and of the new hydrogen ion content
to their in situ values:(a) at 2°C, 1 atm pressure?
(b) at 24°C, 1 atm pressure?
Assume that there is no L:C02, alkalinity or salinity change
during storage.

2-2

Wha t

2-3

Organic tissue containing 0.2 moles of carbon is added to a
1000-liter sample of deep water from the North Pacific bottom
water.
The bacteria in the sample respond and consume th~organic matter drawing on the residual oxygen (100 x 10moles/I).
Which respiration ingredient will run out first oxygen or tissue?
At this ingredient's point of disappearance, by how much will the carbonate ion concentration havechanged?
Again use the simplifying assumptions adopted in
the text.

2-4

If the rate of sea floor subsidence averages 30 meters per
million years, what is the age of the sea floor when it passes through the calci te lysocline?
(Assume the ridge crest
elevation to be 2500 meters and the calcite lysocline to beat 4000 meters.) If the mean rate of accumulation of noncarbonate sediment is 0.5 cm/10 3 yrs, how deep would one have tobore into the sea floor to reach CaC0 3-bearing sediment at awater depth of 5000 meters?
If the mean CaC03 accumulatioq
rate at depths shallower than the lysocline is 1.5 cm/10
yrs, now much further would you have to bore to reach thebasalt at the base of the CaC03-bearing sediment?

2-5

In a given area of the sea floor the rain rate of opal is 1.0
g/cm 2 10 3 yrs, that of calcite is 1.0 g/cm 2 10 3 yrs, and thatof alumino silicates 0.2 g/cm 2 10 3 yrs.
The potential
dissolution rate of opal is 1 g/cm 2 10 3 yrs and that or
calcite is 2 g/cm 2 10 yrs for this geographic location andwater depth.
If the respective solution rates depend on thesquare root of the opal and calcite content of the sediment
what will be the steady state opal and calcite content of thesediment?

is the alkalini ty of each of the following solu tions?
(a) One mole of NaCI dissolved in 400 liters of C02-free
distilled water.
(b) One mole of Na2C03 dissolved in 1000 liters of
CO 2-free distilled water.
(c) One mole of CaS04 dissolved in 5000 liters of
CO 2-free distilled water.
(d) One mole of CaS04 and 2 moles of Ca(HC03)2 dissolvedin 3000 liters of CO 2-free distilled water.

107

2-6

A Mn nodule grows at a radial ra te of 2 mm/10 6 years and is
20 percent manganese by weight.
The underlying sediment isaccumulating at the rate of 0.2 cm/10 3 yrs and contains
11,000 ppm of manganese.
Per unit area (as viewed fromabove), which is "capturing" more fm - the nodule or theadjacent sediment?
What is the ratio of these accumulation
rates?
Assume the sediment to have 0.6 g of mineral matterper cm 3 and the nodule 2.5 g of Mn plus Fe oxides per cm 3 •The Mn to Fe ratio in the nodules is two.

2-7

a frequency of 10/m 2 on a gi ven region of the sea floor.
They average 8 centimeters in diameter, and each has a shark's tooth at its center.
If these
nodules grow at the rate of 4 mm/10 6 yrs and if all the teeth
from all the sharks living in the overlying surface water end
up in nodules, how many sharks would you expect to find living in each square kilometer of surface water?
(The average
lifetime of a shark is 20 years; each shark produces 200
teeth during its lifetime.)
Mn nodules are found at

SUPERPROBLEM #2

A sedimentologist named Taco Hashish, fascinated with the
problem of what controls the proportions of opal, calcite, andalumino-silicates in deep sea sediments constructs a large tank inwhich to conduct solution experiments.
His tank has a stirring
device which not only keeps the water well mixed but also disperses any sediment grains that he adds so that they settle uniformly over the bottom.
Hashish fits the bottom with a set ofrotating "spits" designed to homogenize the sediment which accumulates.
Finally, he installs a pump which injects new water into
the tank.
The amount of water in the tank is kept constant by an
overflow.
He then chooses three types of sediment grains.
The
first is a silicate known for its insolubili ty in water.
The
second is a sparsely soluble organic compound which he terms calcilite.
The third is a moderately soluble organic compound which
he terms opali te.
Before starting his experiments Hashish measures the solubili ty of calcili te and opali te.
He also devises ameans of measuring the pore water resaturation time of calcilite
and opalite.
His results are as follows:
solUbility
gm/liter
calcilite
opalitesaturation timesec

2.0 x 10- 3
1.0 x 10- 1

Hashish then starts his first experiment.
He fills his tank withwater free of dissolved opalite and calcilite and pumps in thesame water at the rate of 1 liter/minute.
(His tank has a volumeof 1000 liters.)
Hashish then starts to inject his three crystal
types at the ra te of 1 x 10- 3 g/minu te each.
The grains become108dispersed in the water but settle to the bottom a few seconds
after their injection (the area of the tank bottom is 1 square meter) .
He then lets the pump and the grain injector run until 6
cm of sediment has accumulated. *He then installs a device
which removes sediment as fast as it accumulates (i.e., the thickness of the sediment remains the same).
After running the system
for some time he stops it and removes the water and performs achemical analysis of the accumulated sediment to find the proportions of calcilite, opalite and silicate. He also analyzes a sample of the wa ter for the amount of dissol ved calcili te and theamount of dissolved opalite.
Assuming that all the solution takes
place in the pores, and that upon dissol ving opali te and calcilite, produce a single neutral ion rather than a pair of oppositely-charged ions, predict the steady state water and sediment
composition (the molecular diffusivity of each dissolved productin the sedimen~ is 2 x 10- 6 cm 2 /sec). What would the T values for
calcilite and opalite be?
How would the pore water profile look?
How thick a "stagnant film" would it take to void these calculations?
In what proportions would Hashish have to add the grains
if he wished to get proportions of 8 parts silicate, 1 part calcilite and 1 part opalite in the accumulated sediment?
*Fortunately Taco is a patient aI)d long-lived man.
If the densityof these mineral phases is 2 g/cm and the porosity of the sediment4
is 50%, this sediment would weigh 6 x 10 grams.
Depending on the
ext~nt of solution the accumulation rate would be between 1 and 3 x
710grams per minute.
Hence about 4 x 10
minutes (i.e., 75
years) would be required for the sediment column to reach its
steady state thickness of 6 cm.

LAST MINUTE ADDITION (SEPTEMBER 1982)
A new method for the study of the extent of CaCD 3 solution
within the pore waters of sediments has been independently proposed by Curray and Sayles at Woods Hole Oceanographic and Quayand Emerson at the Uni versi ty of Washington.
Both groups havemeasurements of the o13 C and LC0 2 in sediment pore waters.
These
profiles provide proof that significant CaCO 3 solution is occurring.
Honjo and his coworkers at Woods Hole Oceanographic have usedtheir in situ calcite exposure method to study opal (720).
Their
results show a dramatic temperature dependence of the solution
rate.
It appears to be about 100 times greater at 25°C than at
2°C!

109

THE ATMOSPHERIC IMPRINT
THE CYCLES OF GAS WITHIN THE SEA

INTRODUCTION
Overlying and interacting with the ocean is the atmosphere.
In our discussion in chapter 1 of the chemical inhomogenei ties
within the ocean, we did not treat the dissolved gases because
their distribution is influenced not only by the formation and thedestruction of particulate material but also by their interchange
wi th the atmosphere.
It is now time to discuss this interchange
process and its influence on marine chemistry.
The elements listed in table 1-1 include a series of six thatare chemically inert - helium (He), neon (Ne), argon (Ar), krypton(Kr), xenon (Xe), and radon (Rn).
As we shall see, isotopes offour of these gases (3 He , 39 Ar , 85 Kr , and 222 Rn ), have important
applications as tracers in the sea.
Nitrogen (N 2), although it isconverted to and from nitrate ion (N03-) by various organisms
living in soils and in natural waters, is much like the six inert
gases.
It does not undergo any chemical reactions which measurably alter its abundance in sea water.
In addition to these inert gases, two other dissolved gases oxygen and carbon dioxide - are of particular interest to us here.
These two gases are, as we already have seen, heavily invol ved
wi th the chemical processes taking place wi thin the sea.
Hence
their dis tribu tions will occupy much of our attention in thischapter.
Oceanographers have long known that oxygen gas is, as
the resul t of consumption by animals and bac teria, deficient inmost subsurface ocean water.
The longer a water mass is isolated
from the atmosphere, the lower its oxygen content becomes.
Thisfact has led to the wide use of 02 as a circulation tracer.
Carbon dioxide (CO 2) has a characteristic not shared by the othergases.
It reacts wi th CO 3= ion to produce two RC03- ions.
Insurface water, on the average only about one carbon atom in 200 isin the form of C02 gas.
Because of its large reservoir of HC03and C0 3 = the ocean ultimately determines the atmosphere's C02
content, rather than vice versa.*
There is one other gas we shall consider.
During respiration
not all the nitrogen bound in organic debris is released as
nitrate ion. A small fraction becomes nitrous oxide (i.e., N20).
*This is not true for the other gases.
Since over 95 percent of
the total of the inventory of these gases is found in theatmosphere, the atmosphere dictates the ocean's gas content.

110

The distribution of this gas in the sea provides some clues
regarding the cycle of nitrogen in the sea.
The amount of all the dissolved gases in surface ocean water
is controlled by their concentrations in the overlying atmosphereand by temperature. * The task of establishing these surface
ocean concentrations is simplified by the fact that the atmosphereis almost perfectly mixed with regard to the gases we will examine
here.
Even for C02 (the most likely to vary), the difference inthe atmospheric concentration from one place to another is only afew percent.
The assumption that the air over every region ofocean has the same composition proves quite adequate for our purposes.
The total pressure the atmosphere exerts on the sea surface is also very nearly the same from place to place.
It varies
only because of the highs and lows of the meteorologist.
All deep water was once at the sea surface.
When this water
lost contact with the atmosphere, it carried with it a near equilibrium amount of each atmospheric gas.
So when we examine deepsea water, we would expect the content of its gases (except C02,
02, He and, of course, Rn) to be determined by their atmosphericcontent and by the temperature of the water when it left the surface.
This has been checked and, indeed, the amount of each ofthese gases in deep water proves to be within a few percent of the
expected amount.
The deep waters bear witness to the fact thatthey were once equilibrated with air.**

SOLUBILITIES OF GASES IN SEA WATER
Let us first consider how the solubilities of gases vary withmolecular weight and wi th temperature.
The data in table 3-1
portray a correlation between the solubilities of the noble gases
and their molecular weights.
Note that the higher the mass of thenoble gas, the more soluble it is in sea water. Helium (mass = 4)
*Radon is an exception.
As this gas 1S radioactive with ahalf-life of only four days, its partial pressure in sea and air
never become equal.
As we shall see, the concentration of thisisotope in the surface ocean is of great importance to our understanding of the rate of flow of gases between the air and sea.
**In making this comparison, the potential temperature of thewater rather than t he in situ temperature mus t be u sed.
The pressure on a parcel of water increases from 1 to about 500 atmospheres-as it sinks from the surface of the ocean to a water depthof 5000 meters.
This pressure increase resul ts in a slight increase in the density of the water, which in turn leads to a risein its temperature.
For a 5000 meter descent, this warming averages about 0.4°C.
Thus, thermometers sent to the abyss record
temperatures higher than the temperature of the water when it began its descent [When comparing the propert ies of various waters
within the deep sea, oceanographers use a temperature which has
been corrected for the pressure-induced warming rather than theactual temperature measured using an in situ thermometer.
This
corrected temperature is called the potential temperature of thewater.

111

Table 3-1. Solubilities of various gases in sea water per atmosphere pressureof that gas* . .

.........

.........
N

I

I

I

Gas

Molecular
Weightg/mole

He
Ne
N2
02
Ar
Kr
Xe
Rn

4
20
28
32
40
84
131
222

CO2
N 20
CC13F

44
44
137

I

Solubility
cc/liter

Solubility**
10- 3 moles/kg

I S~~~~~lf~y I

VolumePartition
Coefficient(Air/Wa ter)at 24° C

I

I
Ref.

0° C

24° C

0° C

7.8
10.1
18.3
38.7
42.1
85.6
192
406

7.4
8.6
11.8
23.7
26.0
46.2
99
186

0.34
0.44
0.80
1.69
1.83
3.8
8.4
17 . 8

0.32
0.37
0.51
1.03
1.13
2.1
4.3
8.1

1.1
1.2
1.6
1.6
1.6
1.8
1·9
2.2

135
116
85
42
39
22
10
5.4

(655 )
(655)
(653)

666
476
677

63
47

29
21
29

2.2
2.2
-

1.5
2.1
1.5

( 656)(659 )
( + )

1437
1071

-

-

°

24° C

(653)

(654)
( 658)(660)(660 )*The solubili ties of He, Ne,
2, N 2, Ar and CO 2 in sea water are those recommended by
Weiss of Scripps Insti tu tion of Oceanography.
The solubili ties of Xe and Rn are fromthe CRC Handbook of Chemistry and Physics wi th the salt effect extrapolated from othergases.
**Average density of sea water = 1.025 g/cm 3+Personal communication P.W. Balls and P.S. Liss, East Anglia Univ., Norwich, England.

is the least soluble; radon (mass = 222) is the most soluble ofthose .listed.
The situation for multiatom reactive gases is more
complicated because of their chemical reactivity.
In the third and fourth columns of the table, the solubilit ief? are gi ven as the numbe r of standard cc I s of gas * which
would dissolve at OoC and at 24°c in 1 liter of sea water at 1
atmosphere pressure of that gas.
For example, if we were to
expose sea water at OoC to pure helium gas at a pressure of 1 atm,
8 standard cc's of He would dissolve in each liter of water; for 1
atm of xenon, 192 standard cc ' s would dissolve per liter.
In thefifth and sixth columns the solubilities are restated in units ofmoles/kg.
In the last column of table 3-1 the solubilities are express ed as volume parti tion coefficients.
Think of the following
experiment.
One Ii ter of sea water is placed in contact wi th 1liter of air.
Some xenon is injected into this system, andallowed to distribute between the water and gas phases at 24°C.
It would turn out that for every atom of Xe that takes up residence in the water 10 atoms of Xe would take up residence in theair.
The volume distribution coefficient of xenon is 10 (at24°C) •
If we were to repeat this experiment wi th helium, forevery He atom found in the water 135 would be found in the air.
The volume distribution coefficient of helium is 135 (at 24°C).
All gases become more soluble in water as the temperature isdecreased.
Table 3-1 also lists the solubilities of a number ofgases in sea wa ter at OoC and 24°C. Figure 3-1 presents this same
information graphically.
Note that the temperature dependence of
He is small, changing by only 5 percent over the 24°C temperature
range, whereas the temperature dependence of Xe is quite large,changing by a factor of 2 over the same temperature interval.
Table 3- 2 gi ves the composi tion of wa ter- vapor- saturated ai r
and the corresponding amounts of each gas to be expected in surface sea water at OoC and at 24°C (obtained by mul tiplying thefraction of each gas in air times its solubility at 1 atmpressure) .

THE RATE OF GAS EXCHANGE
By making a rather simple model we can establish the magnitude of the exchange flux of gases between the ocean surface andthe atmosphere.
Our model assumes that the upper few meters of
the ocean have a uniform concentration of the gas of interest.
The same assumption is made for the column of air above the sea
surface.
These two well-mixed reservoirs are assumed to be separated from one another by a "stagnant" film of water. Gases crossthis boundary layer only by molecular diffusion (the random motionof individual gas molecules).
Although diffusive processes are
nondirectional, if a concentration gradient exists, this chaotic
motion will lead to a net transfer of material from the zoneofhighconcentration
to
the
zone
of
low
concentration.
*One standard cc of a gas 1S the amount contained at 1 atmpressure and O°C temperature in a volume of one cubic centimeter.
The units cc and ml are interchangeable.

113

Table 3-2. Equilibrium concentrations of various gases in surface
ocean water.*

Gas

..-

..-

~

PartialPressure
in Dry Air
atm

Ar
Kr
Xe

5.2 x
1.8 x
0.781
0.209
9 •3 x1. 1 x
8.6 x

CO 2
N 20

3.2 x 10- 4
3.0 X 10- 7

He
Ne
N2
02

10- 6
10- 5
10- 3
10- 6
10- 8

Equilibrium Concentrationin Surface Sea Water
cc/litermoles/kg
OOC
aoc
24°c
4.1 x
1.8 X
14.3
8.1
o. 39
9.4 X
1.7 x

10- 5
10- 4

10- 5
10- 5

0.46
3.2 X 10- 4

3.8 x
1.5 X
9.2
5.0
0.24
5.1 x
8.5 x

10- 5
10- 4

10- 5
10- 6

0.21
1.4 X 10- 4*Salinity is assumed to be 35.00%0.

1.8
7.9
6.2
3.5
1.7
4.2
7.2

X

x
x
x
x
x
X

10- 9
10- 9
10- 4
10- 4
10- 5
10- 9
10- 10

2.0 x 10- 5
1. 4 x 10- 8

1.7
6. 7
4.0
2.2
1.1
2.3
3.7

24°C

x 10- 9x 10- 9x 10- 410- 4
x 10- 5
9
X 1010X 10X

9.3 X 10- 6
6. 3 x 10- 9

100 ....-.......- -....--r-----,.--..,..---t90,.....,

80

E
....0

70

"-

E

'-"

>-

60

J-

--.J 50-(IJ::::::>

-I

40

0

(f)

(f)

30

<t

~

20
N2

Ne

10

He

0

0

5

10

15

20
25
TEMPERATU RE (Oe)30

Figure 3-1. The solubili ties of various gases in sea water as afunction of temperature.
The units are standard cubic centimetersof gas contained by a liter of water per atmosphere of pressure
exerted by the gas.

115

Thus, if the concentration of gas in the water is not at equilibriwn wi th the concentra tion of gas in the air, there will be a
net flow of gas through the stagnant film.
The rate at which thegas is transferred across the air-sea interface then depends upon:
1. the thickness of the film (the thicker the film, thelonger the gas molecules will take to wander through it);
2. the rate at which the gas molcules diffuse through sea
water (the warmer the water, the more rapid the molecularmotion) ;3. the magnitude of the disequilibrium between the gas
content of the air and the gas content of the water (thegreater the concentration gradient, the more rapid the
diffusive transfer).
Figure 3-2 shows the concentra tion of the gas as a func tionof distance from the interface, as depicted by our model.
Beneaththe base of the stagnant film, the concentration is constant andequal to that in the surface sea water.
At the very top of thestagnant film, the concentration is at equilibrium with theoverlying air:the uppermost molecular layers of the film areassumed to be in rapid contact wi th the air and to have a gas
content exactly equal to that dictated by the pressure of the gas
in the air and by the air tempe ra ture.
Be tween the top and thebottom of the boundary film, diffusion creates a linear gradationfrom the atmosphere-dictated value to the ocean-dictated value.
If the partial pressure of the gas in the air happened toyield a concentration of gas at the top of the film equal to theconcentra tion below the film, then equilibrium would exist andthere would be no net transfer of gas.
As much gas would enter
the sea as would leave it.
If, on the other hand, the gas
concentration at the top of the boundary layer was higher or lower
than the gas concentration in the water, then a gradient would beestablished and molecules of gas would be transferred through thestagnant film.
The rate of this transfer, or the flux of the gas,
is s ta ted as the amount of gas going into or out of a given uni tof area of sea surface during a given unit of time.
We will stategas fluxes in uni ts of moles of gas transferred across one square
meter in one year (moles/m 2 yr).
The flux of any diffusing substance is proportional to theconcentration gradient.
The proportionali ty constant is referred
to as the coefficient of molecular diffusion.
The concentra tion
gradient is obtained by subtracting the concentration of the gas
at the base of the boundary layer from the concentra tion of thegas at the" top of the boundary layer and dividing by the thicknessof the layer. Therefore:

F

==

D

[gasJtop - [gasJbottom
zfilm

3-1

where F is the gas flux, [gasJ top and [gasJbottorn are the gas
concentrations at the top and the base of the stagnant boundary
layer, zfilm is the thickness of the boundary layer, and D isthe coefficient of molecular diffusion.
Ions in solution diffuse
at room temperature at a rate of about 2 x 10- 5 cm 2 /sec or ap-

116

Gas Content

Air

•

J

Film

EqUi'I'b'
f num
with
Overlyingof Water

Air

Stagnant

SurfaceWater

surface/
Ocean
Concen tra tion

Figure 3-2. Gas exchange model.
A thin film of TIs tagnant Tl water
separates the well mixed overlying air from the well mixed underlying surface water.
Gases are transferred between air and water
only by molecular diffusion through this film.
The concentrationof gas within the film grades from that corresponding to equilibrium wi th the overlying air at the top to that found in the surface ocean at the bas e.
The film thickness decreas es as the degree of agitation of the interface increases.
In the ocean, thisfilm averages about 40 microns in thickness (a micron is one-millionth of a meter).

117

proximately 6 x lO-~2/yr.
The actual diffusion rates depend on
the tempera ture and on the gas:
D increases wi th tempera ture anddecreases wi th the mass of the diffusing molecule (see table3- 3) .
The thickness of the boundary film depends on the windstress; the higher the wind velocity, the thinner the film.
Whenconcentration difference is placed across a thinner film, theresul ting concentra tion gradient is larger.
Therefore gas flows
back and forth between the air and sea more rapidly during periodsof high winds than during calm periods.
The stagnant film model is not the only one which has beenused to characterize the transfer of gases across the air-waterinterface.
The film replacement model provides another suchcharacterization.
In this model slabs of fluid are assumed toreach the interface and remain for a set time before being
replaced by another slab.
Each slab carries wi th it to thesurface properties identical to those for the bulk fluid.
Whenthe slab reaches the surface its upper few molecular layers
immediately reach equilibrium with the air (as we assumed for the
top of the stagnant film) •Molecular diffusion carries thissurface anomaly into the body of the slab.
The thickness of theslab is assumed to be large enough tha t the concentra tion at its
base remains cons tant during its time of residence at the surface(i.e. the slab can be treated mathematically as a semi-infinite
half space).
When the film is pushed back into the interior of
the fluid it carries with it the excess (or deficiency) of gas it
received during its residence at the surface.
For this model the rate of transfer of gas between air andwater depends on the frequency at which the film is replaced.
The
shorter the residence time of the film at the surface the more
rapidly the gas is transferred.
Thus T, the replacement time,plays the role of z, the stagnant film thickness.
The main difference between the models lies in the dependence of gas exchange
ra te on molecular diffusivi ty.
In the case of the stagnant filmmodel the ra te of gas exchange is proportional to the first powerof molecular diffusivi ty.
In the case of the film replacement
model the ra te of gas exchange is proportional to the square rootof the molecular diffusivity.
Thus if we were to use radon data
to calibra te one model or the other the exchange ra tes pred ic tedfor oxygen by the two models would differ by the square root of
the ratio of the diffusivities of these two gases.
Both theoretical considerations and wind tunnel experiments
involving. the simul taneous measurement of heat and gas transportfavor the film replacement model(lg4).
Field studies involving
the simul taneous measurement of gases differing in d iffusivi ty
support the stagnant film model(179).
Fortunately the differencesin diffusivity among the gases of interest are sufficiently smallthat the choice of models is not critical to any current ocean
studies.

STAGNANT FILM THICKNESS DERIVED FROM NATURAL RADIOCARBON
The 14 C distribution between the atmosphere and the surface
ocean can be used to estimate the average value of the stagnant
boundary layer thickness for the global sea surface.
As we will118

Table 3-3. Rates of molecular diffusion of various gases in sea

water. Also given are the products of the solubility of the gas,
S, and its molecular diffusivity, D, relative to this product foroxygen. This index is useful in evaluating the relative importance of bubble entrainment for various gases.
Gas

Molecular
Weight,g/mole

He
Ne
N2
O2
Ar
Kr
Xe
Rn

4
20
28
32
40
84
131
222

CO 2
N 20

44
44

Diffusion Coefficient*x 10- 5 em 2 / sec(SxD)gas
24° C
0° C(SxD)02

Ref.

2.0
1. 4
1.1
1.2
0.8
0.7
0.7
0.7

4.0
2.8
2.1
2.3
1.5
1. 4
1.4
1. 4

0.54
0.44
o. 45
1.00
0.72
1.2
2.5
4.8

(662,665)(662,665)(662,665)(662,665)(662,665)(664,663a)(664,663a)(661,662
666)

1.0
1.0

1.9
2.0

23
18

(662,663)(662,663)*A wide range of experimental results are ~iven in the literature.
Discussion of the possible causes for this variability is beyond
the scope of this book. Measurements in sea water have yet to be
carried out to confirm these results. The temperature effect isderived from reference 662. The values for Kr and Xe are calculated using diffusivity ratios given in reference 664.

119

discuss in more detail in chapter 5, 14 C atoms are born in theatmosphere, where they become 14CO 2 molecules.
They are transferred into the ocean by gas exchange.
By studying the exchange
pattern of this radioacti ve breed of carbon dioxide, we can findthe exchange rates of C02 and other gases.
Prior to anthropogenic alterations the 14C. cycle has beenrunning for so long a time that it should have been at a steadystate; in other words, the amount of 14C. at any point in thesystem should have been constantwith time.
If this is true,
then the number of 14C atoms that were added to the ocean eachyear by gas exchange with the atmosphere must have exactly
balanced the number of 14C atoms that disappeared within the ocean
by radioactive decay.
To determine how many 14C atoms enter theocean every year as a result of gas exchange, we multiply theconcentration gradient of 14 C02 arross the stagnant boundary layer
by the diffusion coefficient for 4C02 • This can be stated in theform of an equation as:
Input of14[[ 14 C02 J top - [14 C02 J bottom
C = D
zfilm

J Aocean 3-2

14C is measured in ratio to normal carbon, or C*/C.
It is
therefore appropriate to reexpress the 14 C02 concentration as:

3-3
and

3-4

where [C02Jbottom and [C0 2 Jtop are the CO 2 gas concentrationsa t the base and the top of the stagnant boundary layer, and theratio a14C02/aC02 reflects the fact that because of their greater mass 14 C02 molecules (molecular weight = 46) are 1.5 percent
more soluble than ordinary C02 molecules (molecular weight = 44).
Thus the 14C/ C ra tio at the to~ of the boundary film is sligh tly
greater (1.7 percent) than the 4C/ C ratio measured in atmosphericcarbon.
At steady state there can be no net transfer
carbon between the ocean and atmosphere. Hence:[C02Jtop= [ C0 2 J bottom = [Co 2 J sur face
If so,
as:equation 3-2 for the input of

In~~g ofD[ CO 2] s urfaceAocean

Zfilm

[a

14C atomsofordinary

3-5

can be rewritten

3-6
14 C0 2
aC0 2

120

C*)
C atm

C*)
]
C surfaceThe number of 14 C atoms being lost by radiodecay must match
this gain.
To find an ex~ression for the loss rate per year, we
begin with the number of 1 C atoms in the sea. To calculate this,
we multiply the volume of the ocean (expressed as the area of thesurface ocean, Aocean, times the mean depth of the ocean, h) by
the average total dissolved carbon content and then by the averagera tio of 14C/ C in this oceanic carbon.
We then mul tiply thisnumber of 14 C atoms in the sea by the fraction undergoing radioactive decay per year (i.e., by x). This yields:
Loss of 14 C = Aoceanh [LC02Jocean

C~ A
C hcean

3-7

Setting the right hand sides of equations 3-6 and 3-7 equal
to one another and solving for zfilm, we obtain:

zfilm

D [C02Jsurface
Xh [LC02Jocean

1 -(C*/C)surface

aC02(C*/C)atm(C */C ) ocean

a14C02(C*/C)atm

a14C02

3-8

exc02

The average value for the molecular diffusion coefficient, D,
is 5 X 10- 2 m2/yr for C02 in surface sea water, the mean depth ofthe ocean, h, is 3800 meters, and the reciprocal of the decay constant of 14C, 1/ A, is 8200 years.
The average concentration of CO 2
gas in the surface ocean is 0.012 moles/m 3 , the average LC02 content of ocean wa te r is 2.2 moles/m 3, and the aco 2/ a 14CO 2 ratio
is 0.983.
The 14C/C ra tio in average ocean carbon was, prior to
nuclear tes ting, about 0.83 that in atmospheric CO 2.
The 14C/ C
ratio in surface ocean carbon is, on the average, 4 percent lower
than the 14C/ C ratio in atmospheric carbon dioxide.
This lastfactor is the one that leads to the largest uncertainty in thedetermination of zfilm.
The difference is known to no better
than one ~ercent.
A one percent error in the measurement of theaverage 1 C/C ratio in surface water carbon relative to that in
atmospheric carbon would create a 15 percent error in the estimateof zfilmI
When zfilm is calculated from these figures we obtain 40 x
10- 6 meters. For lengths this small, it is more convenient to use
the unit -micron (one-millionth of a meter).
So zfilm is 40
microns.
Keep in mind that the boundary should be thinner than
this average value in areas of high wind velocity and thicker inareas of low wind velocity; our calculation gives us only an
oceanwide average.
Since the coefficients of molecular diffusion do not vary
greatly from gas to gas, it proves convenient to think of the gas
exchange process in terms of the ratio D/z.
Because thedimensions of D and zfilm are, respectively, m2 /yr and meters,
the ratio D/z has the dimension of m/yr (hence those of velocity).
We might think of this ratio as the velocity of two
imaginary pistons:
one moving up through the water pushing ahead121

of it a column of gas with the concentration in surface water, andone moving down into the sea carrying a column of gas wi th theatmosphere equilibrium concentration found in the upper few
molecular layers of the water column.
On the average thes
2 2
pistons push C02 at the rate of 5 x 10- m /yr divided by 40 x 10meters or 1200 m/yr.
Thus each day a column of sea water about
3.3 meters thick exchanges its gas with the atmosphere!
The
piston concept clarifies the fact that even if the atmosphere andocean are at equilibrium, the transfer of gas continues:
theamount "pushed in" just matches the amount "pushed out."
From the piston velocity we can calculate another importantgeochemical time constant, i.e., the characteristic time requiredfor ordinary gases in the ocean surface mixed layer to equilibratewith the atmosphere.
Since the mixed layer ranges in thickness
from 20 to 100 meters with a piston velocity of 3 meters a day theequilibration time lies in the range of one week to one month.
Because of their differing diffusivities the flushing times forvarious gases will differ somewhat from one another.
These differences are summarized in table 3-4.
02 has a diffusi vi ty (andhence a piston velocity) about 20% higher than that of C02.
For
Rn the diffusivity (and hence piston velocity) is about 25% lower
than that for CD 2.
As shown in this table for any gi ven film
thickness the piston velocity will also change with temperature(because the diffusivity changes). However, because of the change
in solubility with temperature the absolute invasion rate changes
more slowly with temperature than does the piston velocity.

t

STAGNANT FILM THICKNESSES AS DETERMINED BY THE RADON METHOD
Another tracer method exists which allows the rate of exchange in local areas at the ocean to be measured.
It involves
the escape of the radioactive gas 222 Rn from the sea to theoverlying air.
Radon is generated in sea water by the decay of
its parent 226 Ra .
Away from the interfaces with air and sediment
the radioactivity of radon is identical to that of its parent.
Howeve r, as we shall dis cuss in chapter 7, radon leaks into thesea from its sediments, giving bottom waters a higher radon
acti vi ty than expected from the radium content of the water.
Of
interest in this chapter is the fact that the surface layer of theocean has a deficiency of radon activity. This is because the air
over the open ocean is nearly devoid of radon.
Hence the flux ofradon atoms from sea to air is not matched by a return flux.
Someof the raQon atoms generated in the surface layer of the sea
escape to the atmosphere before undergoing radiodecay.
The fraction which escapes provides a means for determining the piston
velocity of gases as a function of wind velocity.
The task of finding the deficiency of radon atoms in thesurface water column is greatly simplified by the fact that inmost regions of the surface ocean there is a wind-mixed layer from
20- to 100-meters thick in which radon atoms are well mixed. This
layer is underlain by the upper thermocline which mixes verticallyat such a very slow rate that none of the radon atoms in this zone
reach the overlying water during their average lifetime of 5.5
days.
This being the case, the flux, F, of radon atoms from a122Table 3-4. Invasion rates and equilibration times of various gases assuming astagnant film thickness of 40 microns and a mixed layer depth of 40 meters~
24 0 c

oOC
D

Gas

.f'.)
c:,A,)

He
N2

02
Rn
C02
N2 0

v

T

m

days

10- 5 cm 2
sec

day

2.0
1.1
1.2
0.7
1.0
1.0

4.3
2.4
2.6
1.5
2.2
2.2

9
17
15
27

18* *

Ie

S

I

moles

moles

m3

m 2 yr

1.8xIO- 6 2.8xl0- 3
0.64
561
0.36
342
0.021 I
17
1.4xI0- 5 1.1x10- 2

D
10- 5cm 2
sec

4.0
2.1
2.3
1.4
1.9
2.0

v
m

T

I

m3
1.7x10- 9
0.41
0.23

5.3x10673
417

- 3
13
10* * 9.5x10- 6
6.5x109142
1.0x10-

day

8.6

5

4.5
5.0
3.0
4.1

9
8

4.3

S
moles

moles
m 2 yrdays

3*D stands for the molecular diffusivity, v for the piston velocity, T for themixed layer equilibration time, S for the equilibrium concentration in surface
water and I for the invasion rate of the gas into the sea.
**Neglecting the chemical influence of HC03- and C03= (i.e., for a Coca Cola ocean).

square meter of sea surface to the atmosphere can be estimated asfollows:

3-9

where h is the thickness of the wind-mixed layer at the ocean surface,
A2 26
and A222~n' the respective decay constants for
226 Ra and !\ft Rn , [226Ra:~ the concentration of radium in the surface-mixed layer and [22 Rn] the concentration of radon in surface
water.
As the product decay constant times the concentration isthe activity, A, of the respective isotopes, this equation can be
rewritten:
F = h (A226 Ra - A222 Rn )3-10
Using the stagnant film model, the flux of radon atoms canalso be written as:
3-11
or

F

=

D222 Rn A222 Rn
zA 222 Rn

3-12

Setting these two expressions for radon flux equal to one another(i.e., assuming that steady state exists for the radon concentration in surface water) and solving for z, we get:z =D222 Rn
A222 Rnh1

3-13

An example of the application of this method is given in figure 3-3.
Beneath the mixed layer (as defined by the temperature
profile) the radon activity is found to be within the measurement
error of that of radium.
By contrast, within the mixed layer theradium activity is about 1.6 times the radon activity.
The mixed
layer thickness in this example is 40 meters, the decay constantfor radon is 2.1 x 10- 6sec- 1 , and the diffusion rate of radon 1.4
x 10- 9 m2/sec.
Hence:z =92

1.4x10- m /sec
2.1x10- 6 sec- 1 x 40m

[1]
1.6-1

28x10- 6 meters
28 microns
sea.
ized

A large number of such profiles have been determined in the
The stagnant film thicknesses derived from them are summarin the histograms of figure 3-4.
In the north and south124

Rn ACTIVITY (dpm/IOOI)
8
3
4
5
6
7
2
O;....--w.--.......--...~..,..,.~,.",..~-r-r-----:-: .....,

·····
···
···
··
··-

20

·

40

E
::I:
I-

60

a..

W

o 80

.........

...

...
...
..

..

..
.. ...

_.0.
.. -..
.-.

.0100
I 20 ......._ - - I_ _--.l._ _.-a18
19
20
21

....-..-.-_----'

22

23

24

TEMPERATURE (Oe)Figure 3-3; Concentra tion of radon gas as a function of depth atAtlantic GEOSECS sta tion 57 (24° S 35° W) •The concentra tion expected if no radon were escaping to the atmosphere is shown by
the dashed line.
The difference between this equilibrium value
and the observed value is a measure of the amount of radon lost to
the atmosphere.
The shaded area is a measure of the total amountof radon lost to the air.
From the amount of missing radon it ispossible to determine the piston veloci ty for gas exchange. The
dotted line is the temperature profile at this station defining
the depth of the wind stirred layer.
These results were obtained
as part of the GEOSECS program (175).

125

PACIFIC

ATLANTIC

o 20 40 60 80 100 120 a 20 40 60 80 100 120

a 20 40 60 80 100 120 a 20 40 60 80 100 120
~

I

I

I

I

I

I

a 20 40 60 80 100 120 a 20 40 60 80 100 120
-

-

I

I

a 20 40 60 80 100 120 a 20 40 60 80 100 120;] (microns);] (microns)

Figure 3-4. Histograms of stagnant film-thickness estimates basedon radon measurements in the surface wa ters of the Atlantic andPacific Oceans.
The measurements are separated according to latitude belt.
There is a suggestion that the exchange rate is somewhat lower than average in the equatorial zone (15°8 to 15°N) andsomewhat higher than average in the Antarctic zone (>50 0 S). These
measurements were made as part of the GEOSECS program (175, 176).

126

temperate zones the values fall largely in the range 20 to 60
microns; in the equa torial zone the values fall mainly in therange 20 to 80 mic rons; in the An tarc tic they are mainly . in therange 10 to 40 microns.
This distribution is consisten~ wlth thefact that wind velocities are on the average somewhat hlgher over
the Antarctic Ocean and somewhat lower over the equatorial ocean
than over the temperate ocean.
However, as none of the high latitude measurements were made in winter (i.e., the period of greattest storminess), the difference between the various zones may beeven larger than suggested by these results:. .
A comparison between stagnant film thlcknesses measured In awind tunnel and those measured by the radon method at sea in areas
with differing mean wind velocities are shown in figure 3-5.
As
can be seen, while the trends are similar the environmental film
measurements are larger at any given wind veloci ty than thosemeasured in wind tunnel s.
~1ore field res earch will have to be
carried out before a firm estimate of the wind velocity dependenceof the gas exchange rate can be given.

OXYGEN CONCENTRATIONS IN SURFACE OCEAN WATER
When the dissolved oxygen contents of surface ocean waters
are compared with those for saturation with the overlying air it
is found that they are on the average supersaturated by several
percent (see figure 3-6).
Since subsurface waters are all highly
undersaturated with dissolved oxygen, this excess is not theresult of incomplete equilibration between upwelled water and theatmosphere.
There are several possible explanations for departures from equilibrium.
1) Bubbles inj ec ted into the wa ter lose part of their gas to
the water during entrainment.
As any excess introduced in thisway must be balanced by diffusion through the stagnant film backto the atmosphere,
bubble entrainment could maintain an overpressure of gases in surface water.
2) Plants living in surface wa ter produce oxygen.
The accumulation of this oxygen could lead to observed supersaturation.
3)
The wa ters on which these measurements were made were
undergoing seasonal warming, and hence had dissolved oxygen
contents reflec ting previous lower tempera tures.
As the GEOSECS
measurements were made in the fall as well as in the spring andsummer, and as in equatorial waters there is no seasonality, this
explanatiop is not acceptable.
Our knowledge of the rate at which dissolved oxygen goes backand forth between the a tmosphere and surface sea helps us todistinguish between possibilities 1) and 2).
Let us firstconsider the possibility that photosynthesis produces the excess.
\Jarm surface water at equilibrium with the air contains about
0.24 moles /m 3 of 02.
Thus a mixed layer 50 meters in thickness2
contains abou t 12 moles/mof 02.
At a pis ton veloci ty of 3
m/day, about 0.7 moles of atmosphere 02 will invade each square
meter of interface each day.
Under these circumstances theaverage 02 molecule is replaced in the ocean-mixed layer once each
12/0.7 or 17 days.
This is the time then that the average 02
molecule produced by photosynthesis will reside in the ocean-127

WIND VELOCITY (knots)
12
16
20
24

28

908i---.-;.p~-~--.,...--.....,r----,---'

o

•

70

•
50

WIND TUNNEL
NATURAL 14C
RADON

SUMMERANTARCTIC

E

)

30

I 0 ' - - - -.....- -......- -......--~-~~ ...
4
6
8
10
12
14
WIND VELOCI TY (m/s)
Figure 3-5. Plot of film thickness against wind velocity.
The solid line represents measurements made in a wind tunnel.
The solid
circles are film thicknesses based on radon measurements in theocean.
The open square represents the global average film thickness based on the distribution of natural radiocarbon.
While boththe open ocean and wind tunnel data show increases in film thickness wi th decreasing wind veloci ty, the tunnel measurements giveconsistently lower film thicknesses than do the ocean measurements.
Th·is offset may reflect the following differences between
the two situations.
The large waves present in the sea are notpresent in the wind tunnel.
Fresh water is used in the wind tunnel experiments.
Also, the method used to adjust the tunnel windveloci ties
(measured a few tens of centimeters above the wa ter
surface) to the equivalent velocities over the sea (measured at aheight of about 15 meters) may not be correct.
The wind tunnel
experiments were conducted by H.C. Broecker and his colleagues in
Hamburg, W. Germany (188).
The radon resul ts were obtained byscientists at the Lamont-Doherty Geological Observatory (172,173,

175,176).

128

SURFACE
WATERS

•

'"'""

Ol

~

"
0

•

E

:l..
--'
C\J

0

0
W

>
.....J--,,-,,:

250

I- _

0

en
en
-

0

200

0

10

20

30

TEMPERATURE (OC)

Figure 3-6. Plot of the dissol ved oxygen content measurements on
surface wa ters made as part of the GEOSECS program as a functionof water temperature (424,425,426).
The solid line represents thesaturation value.
Surface ocean waters are supersaturated by anaverage of 7 ~mol/kg (i.e., by about 3%).

129

mixed layer before escaping to the atmosphere.
The rate of plant
producti vi ty in the ocean-mixed layer has been determined by marine biologists at many points on the sea surface and at manytimes during the year using a tracer 14C~ method developed by aninventi ve Dane named Steeman-Nielsen.
The mean for all of themeasurements made in the open ocean 1s about 0.3 moles of carbon
fixed p~r square meter of ocean surface per 17 days (i.e., about 6
moles/m yr).
As much of the plant matter produced in the sea Is
promptly eaten by animals and bacteria living in the mixed layer,
the rain rate of organic material from the mixed layer is lessthan the production of organic matter within the mixed layer.
Onthe average only about 15% of the matter formed in the mixed layer
falls beneath the mixed layer before being consumed. Thus the net
fixation rate of carbon in the mixed layer is only about 0.045
moles/m 2 per 17 days. For each mole of carbon sinking beneath themixed layer, 1.3 moles of 02 are left behind dissolved in themixed layer.
Thus the net production of oxygen in the mixed layer
is 0.06 moles/m 2 per 17 days.
This may be compared to the atmospheric equilibrium value of 12 moles 02/m2.
The overpressure
needed to exhaust the biogenic oxygen is only 0.06/12 or about
0.5% of total. *Thus production of oxygen by plants does notappear to explain the observed mean overpressure of this gas.
This could only be the case if the amount of organic material
fallin~ fromthe mixed layer were about five times that of 1
mole/m yr estimated here.
There is no doubt that if enough bubbles are entrained and if
these bubbles penetrate to a depth of several meters, the observed oxygen supersatur'ations could be generated.
The pressure
at a depth of two meters is 1.2 atmospheres, hence bubbles carried
to this depth would continue to lose gas to the adjacent water
until a 20% supersaturation was produced.
So the potential isthere.
The difficulty with this hypothesis comes in attempting toquantify the flux of gas borne by bubbles into the sea.
Itdepends in a complex way on the spectrum of winds and waves.
So,
while promising, this hypothesis remains unproven.
One means by which the bubble injection hypothesis could bechecked would be to simul taneously measure the degree of supersaturation of a number of gases.
If the oxygen supersaturation
were produced by photosynthesis then none of the other gases would
show excesses.
If these gases are supersaturated the excesses
should not be the same because the solubili ties of gases differ.
For example, were a bubble to be carried deep enough beneath thesurface ana remain long enough to dissolve completely, the gas it
gave up would produce a much greater percentage of change in thehelium content of the water than in the xenon content of thewater'.
Even if the bubble didn't completely dissolve, thetendency would be the same, and the low-solubility gases would beenhanced in the sea more than the high-solubility gases.
Workingagainst this effect, however, ,""ould be the fact that the lower~e have checked this result using the I-D ocean model presented
1n chapter 10. The model, which yields a reasonable average vertical distribution of P04 in the sea, yields a mean 02 supersaturation for surface water of 0.4%.

130solubili ty gases could escape back through the air-sea interface
fas ter due to their higher molecular diffus i vi ties.
As can beseen in table 3-3 the diffusivity changes almost as much from gas
to gas as does the solubili ty, hence; the solubility effect will
be partly compensated by the diffusivity effect.
In any case,sufficiently accurate measurements for a variety of gases measured
simultaneously in surface sea water are not yet available, so thischeck has yet to be made.
As shown in figure 3-7 a narrow belt of surface ocean along
the equator in the Pacific Ocean shows deficiencies in dissol ved
oxygen.
These deficiencies are maintained by events of upwellingof waters deficient in oxygen.
Because of the intensity of thisprocess, invasion of 02 from the air does not have time to wipeout this deficiency.
OXYGEN DEFIC_IENCIES IN THE DEEP SEA

Even though the oxygen content of surface waters is observed
to be slightly supersaturated, concentration of this gas in newly
formed deep waters is the most predictable of all the biologicallyutilized constituents.
As we have already pointed out, thebiolimiting nutrient constituents (i.e., H4Si04, N03 and P04) areonly partially utilized in cold surface water.
Hence, it isdifficult to predict how much of anyone of these constituentsfound in deep water was produced by respiration within the deepsea and how much was carried down with newly formed deep
water.*The same difficulty applies to ~CD2 and to radiocarbon.
The fact that dissolved oxygen is carried to the sea's
interior in nearly known amounts provides possible means forrecreating the compositions of deep waters at their time oforigin. For any given water in the deep sea, we can determine howmuch oxygen was utilized by subtracting the observed oxygen
content from the saturation oxygen content.
Oceanographers havetermed this quantity the apparent oxygen utilization and refer toit as AGU.
The adjective "apparent" is added to alert users of
this property that it is not an exact measure of utilization.
Thewaters leaving the surface likely carry more than the saturation
amount of oxygen..
Due to the nonlineari ty in the solubili ty ofoxygen wi th temperature, any miXing between waters of differing
temperature will produce a slightly higher oxygen concentration
than would be expected for the intermediate temperature. However,as these effects are small compared to actual utilization, AOU isqUite close to TOU (true oxygen utilization).
Before examining the reliabili ty of AOU in separating the
preformed contribution to the observed deep water concentrationsof the various nutrient constituents, let us look at the disttribution of this property in the sea.
Figure 3-8 shows how AOUchanges along the 4000 meter horizon.
As would be expected, AOU
is lowest in the northwestern Atlantic and highest in the northern
Pacific.
The pattern is similar to that for dissolved silica(see figure 1-8).
However when the vertical sections for AOU*1n the sections which follow we will refer to these concentrations in newly formed deep water as the preformed values.

131

10

~

~

~

"'0

E

:L

......-..

.-<t

en

(\J

0

0

•
1
• • •••.0'•5
OO~•i • m
•
• ••
0
•• •
•
-5
• 8
0

0

0

0

0

0

0

0

(\J

0

<:>

-10

<:>

0

0

t:

0

8

• NOV 1979

SURFACE
WATERS

5

0

5

-15

EQU
LAT I TUDE

S

0

0

•
•
•

o JULY 1979

0N

Figure 3-7. Difference between the observed and saturation 02 contents for surface waters in the equatorial zone between Hawaii andTahiti.
Intense upwelling very close to the equator brings up water deficient in 2 - These measurements were made by Bob Williams
and his group at Scripps Institution of Oceanography as a part ofa program organized by Wyrtki at the University of Hawaii (433).

°

132

60

60

20

20

60

100

140

180

140

100 6040204060---6...0 -........2..0-~...
20~a....-.....
6-0......~10"'0--14""'O~-18""O---14""0----,""O-0""'60

LONG I TUDE

Figure 3-8. Distribution of AGD (saturation oxygen content minus
measured oxygen content) at a depth of 4000 meters.
Based on the
results obtained during the GEOSECS program (424,425,426).

133

(figures 3-9, 3-10, 3-11 and 3-12) are compared with those forsilica (figures 1-14,1-15, 1-16 and 1-17), it can be seen thatthe dis tribu tion of AOU is not the same as tha t of silica.
A
promiment example of this difference can be seen in the wes tern
Atlantic where AOD reaches its highes t value in the equa torial
thermocline while silica reaches its highest value in bottom
waters.
While some of this difference is due to the high preformed s il ica of deep wa ters formed in the Antarc tic, this is not the
only cause.
As mentioned earlier, the frustules of opaline silica
formed by organisms reach, on the average, considerably greater
depths before dissolving than do fragments of organic tissue before being eaten.
Thus, while the distributions of both silica
and AOU reflect the main circulation pattern of waters in the deepsea, they also have important differences generated by processes
occurring at the sea surface (i.e., gas exchange and plant uptake)and wi thin the body of the sea (respira tion and silica solu tion) .
Thus, while AOU may help us unders tand the dis tribu tions of theconstituents of soft tissue (C, Nand P), it will be of no
particular use in unders tanding tha t of hard part consti tuents(Si, Ba and Ca).
The ratio between carbon release and oxygen consumption duringoxidation of organic material has been established in the laboratory.
When the debris of marine plankton is oxidized, abou t 1.3
moles of oxygen gas are required to combust an amount of organic
rna terial containing one mole of carbon.
In other words, in thisexperiment, for each mole of C02 that appears as a combustion product, 1.3 moles of 02 gas disappear. In the reaction the combustion produc ts are CO 2, NO 3-, and H20. .The chemical reac tions are:

3-14

02 + CH20 ~C02 +H2 0
20 2+ NH 3 + OH- ~ NO 3- + 2H 2

°

3-15

where CH 20 and NH3 are the chemical forms of carbon and nitrogenin organic tissue.
About one 02 molecule is required to "burn"each C a tom present in the organic tissue, and about two
2 molecules are required to "burn" each N atom present.
The other constituents which undergo oxidation are so low in abundance relative
to C and N that their 02 requirements can be neglected.
Since
there are about 15 atoms of N for every 105 atoms of C in fallingorganic debris, 135 molecules of 02 are required for each 105
a toms of C.'
While these ratios apply to the average organic matter formedin surface water, they need not necessarily apply to the organic
rna tter falling to the deep sea.
The problem is tha t the respira tion occurring in shallow wa ter may oxid ize one cons ti tuent more
efficiently than another.
If so, the residues falling as fecalpellets need not have the same composition as plants.
Also, insome regions of the deep sea, organisms use N03- ion as an oxidant(they convert N0 3- to N 2 ).
To check this, we show in figures 3-13
and 3-14 the changes in P0 4 versus that in N0 3 , in N0 3 versus thatin 02, in 02 versus that in LC02, and in alkalinity versus that in
LC02 along constant density horizons in the major ocean basins.
We choose constant density surfaces to avoid the problem that deep

°

134

I

IQ..

W

o

2

I

IQ..

W

o

6lL.lU~-L_..L.----I_--l.._.J----IL----L_-'--_.L..---Io_"'"

60
S

40

20

0
LATITUDE

20

40

60
N

Figure 3-9. Vertical section of AOD (saturation oxygen content
minus measured oxygen content) along the western basin of the Atlantic Ocean. The section is broken at 1500 meters. The verticalexaggeration for the upper ocean is twice that for the deep ocean.
The contour interval for the upper ocean is 20 lJIIlol/kg; that forthe deep ocean is 10 ~mol/kg. Also shown by dotted lines are the°0
27.0%0 isopycnal (upper), the 00
27.5%0 isopycnal (middle) and the 04 = 45.2%0 (lower) isopycnal horizons.
Based on the resul ts obtained by the GEOSECS program (424).
The
tracks along which this and the following sections were drawn canbe found on the foldout map.

135

3

3

E

EASTERN
ATLANT IC
AOU
fLmol/kg

E~~........
I

I

l-

I0...

0...

w

w

4

0

0

5

NORTH
ATLANTIC
AOU
fLmol/kg

60
N.AMER.

40

20

LONG ITUDE ow AFRICA

~O
S

o

20

LATITUDE

40
N

Figure 3-10. Vertical sec tions of AOU (sa tura tion oxygen content
minus measured oxygen content) for the waters below 1500 meters
depth across the North Atlantic (at about 32°N) and in the easternbasin of the Atlantic Ocean.
Based on the results obtained by the
GEOSECS program (424).

136

2

..--...

3

E~.........

:c

~4

w
o

5

61-....---L----J1.----JL.----I~---IL...-........~------'

60
S·

40

20
LATITUDE

020
N

Figure 3-11. Vertical section of AGO (saturation oxygen content
minus measured oxygen content) below 1500 meters along the basinsor the western Indian Ocean. Also shown by the dotted line is the
0 4= 45. 2 % 0 isopycnal horizon.
Based on resul ts obtained by
the GEOSECB program (426).

137

I

.-

0W

o

.I

0W

o

6'----1_-L_..J.........I...l..-~~__L._....L__..r.____JL.._........L.

S

60

40

20

0

LATITUDE__=__=_----...

20
N

Figure 3-12. Vertical section of AOU (saturation oxygen concentration minus measured oxygen concentration) along the western
Pacific Ocean.
The section is broken at a depth of 1500 meters.
The vertical exaggeration for the upper ocean is twice that forthe deep sea.
The contour interval for the upper ocean is 25
lJmol/kg.
Tha t for the deep ocean is 10 ~ol/kg.
Also shown by
dotted lines are the 00 = 27.0 % 0 isopycnal (upper), the 00
27.5%0 isopycnal (middle) and the 04 = 45.2 0 /00 (lower)
isopycnal horizons.
Based on results obtained by the GEOSECS
program (425).

138

waters of different orlgln carry different preformed concentrations of P0 4 , N0 3 , EC.Q z and alkalinity.
Within a single ocean
basin, waters on a given density surface have the same origin andhence the same preformed nutrient constituent concentrations.
We
include an alkalinity versus ECOz plot so that the contribution of
C.aC.Q3 solution to the change in the EC.Oz concentration can be
accounted for.
The results of these comparisons are summarized intable 3-5.
The first column in this table summarizes the ratiosof the concentration changes expected if the composi tion of thematerial released during respiration were the same as the composition of average marine plant matter.
As the ratio of CaCD3 toorganic tissue produced depends on ecology rather than biochemistry, no entries are made for those ratios involving ECOz or CaC03.
The upper four entries in the last three columns of this table aretaken directly from the graphs in figures 3-13 and 3-14.
The
lower four entries are calculated from the upper four, using thefollowing relationships:°

L\CaC 3

112

(L\Alk + L\N03)

L\ORG = L\CaC03 -L\ECOz

3-16
3-17

The first of these equations yields the contribution of ECOz derived from CaC0 3 solution.
The second gives the change in ECO z
from the oxidation of organic debris.
The entries in this table make it clear that while the proportions are roughly as expected, a subs tantial range is found.
In particular the ra tio of organic C to PO 4 and to NO 3 for the
Atlantic is smaller than expected.
The ra tio of carbon added todeep water from the oxidation of organic residues to that added by
the dissolution of CaC03 ranges from 2.6 in the Atlantic to 1.3 inthe Indian Ocean. The reason these values are less than the valueof four given in chapter 2 is that the oxidation of organic matteroccurs, on the average, at shallower depths than the solution of
CaC03 (i.e., the ratio L\ORGIL\CaC03 decreases with depth in theocean) .
Thus our strategy of using AOU as a means for reconstructing
the ini tial composi tion of the deep waters in the sea meets wi th
only limited success.
It cannot be used for the constituents
bound in hard parts (this eliminates C as well as Ca, Sa and Si).
Since the composi tion of the debris falling into the deep sea
varies from region to region in the ocean the use of a uni versal
correction. factor will lead to errors in the reconstructed initial
P0 4 and N03 values.

THE MARINE NzO CYCLE
During the oxidation of organic material by organisms livingin the sea some of the nitrogen is converted to nitrous oxide gas(NzO) rather than back to nitrate ion.
Dramatic evidence for thisis shown in figure 3-15.
The profile of dissolved nitrous oxidein the northwestern Atlantic has a shape which is the inverse ofthat for dissolved oxygen!
As shown in the insert in this diagram, when the excess NzO concentration (over that for equilibrium139

Property-Property trends along the 04 = 45.8%0
isopycnal horizon in the deep sea (~2500 meters depth).

Table 3-5Property
Ratio

Reference
Atlantic
Indian
Value
Ocean
Ocean
From Graphs in Figures 3-13 and 3-14

Pacific
Ocean

11
-154
-0.8
0.8

15
-170
-0.9
0.6

~NO 3/ ~P04
~O 2/ ~P04
~O 2/ ~ECO 2
Mlk/ ~EC02

15
-136

-

-

13
-82
-1.0
0.4

Calculated from the above using equations 3-16 and 3-17
~O 2/ ~ORG C~ORG C/ ~NO 3~ORG C/ ~P04
~ORG C/ ~CaCO 3

-1.4
5
59
2.6

-1.3
6.3
105
C -

-1.4
10
110
1.3

-1.4
8
125
1.9

Figure 3-13. The upper diagram on the facing page shows plots ofnitrate versus phosphate concentration and of dissolved oxygen
versusphosphateconcentration along the45.8 0 /00 isopycnal
surface (~ 2500 meter depth) in the Atlantic, Indian and Pacific
Oceans. The slopes of these relationships give the proportions ofthese constituents for respiration taking place in the deep sea.
Based on resul ts obtained by the GEOSECB program (424, 425, 426)
J of
Minster correctly points out that the trends seen for the
deep Atlantic are likely dominated by mixing with Antarctic
waters.
Hence they are not indicative of respiration occurring in
the deep Atlantic.
00Figure 3-14. The lower diagram on the facing page shows plots ofdissolved oxygen versus total dissolved inorganic carbon content
and of alkalinity versus total dissolved inorganic carbon content
along the 4508 0 /00 isopycnal horizon in the Atlantic, Indian andPacific Oceans.
The slopes of these trends reflect the composite
effects of respiration and calcium carbonate solution in the deepsea.
Based on results obtained as part of the GEOSECS program(424, 425, 426).

140

26
~
~

....

24

E

22

'0

:l.

....

20

0

z

I~.O

3.0

280

~

:::::
o

~ 250

230
22 0 '--_I--.lI-.--'_-L_-L_-..L:l~"'"
1.0
1.2
1.4
1.8
PO", (flmollkO)

II 0 - -.......- .......2.0
2.2......._.1.-_.1.-_.1.----:1
2.4
2.6
PO", Cflmollko)

280----------..---,

220r---r----,~-"T"""-~--r--.,

2.4

2.6

2.8

3.0

PO", (p.mollkg)

IeO_~_--_-_._-__.r__-_._-_INDIAN
45.8%.

ATLANTIC
45.8."..

i......
'0
E

5

N

o

240
230

AOI 11.1.0
AlCO!

220 L..._..L._......._--lL.-_...L-_..... 100"--......- - - 1 . - - 1 - - . . 1 - - - ' - - - - 1
2160 2170 2180 2190 2200 2210 2240 2260 2280 2300 2320 2340 2360
IC02 (p. moll1l o1
tC02 (p.rnoIlItOI
2370r--....---,~-.....--r--., 24401---r--......,~-"T"""--...--_r::7i 2470

i

2360

ATLANTIC
45.8."..

INDIAN

i

5 23

...>-

r

~

J

2340

2370 ....--L.--JL-.-..I----L--I---I'
2360 2280
2300 2320 2340 2360 2380 2400
IC02 (p.mollkg)

4~~......

2300 2320 2340 2360 2380 2400
l:C02 (p.mollltg)

PACIFIC
45.8%.

Q

8'\

z

;i

AAlK '080
AlCO!
.

~

<l:

2320L...--L.---I1-.-..I---L---I
2160 2170 2180 2190 2200
IC02 (flrnollkg)

2260

2280 2300 2320
IC02 (flmol/Itg)141

with the atmosphere) is plotted against the deficiency of oxygen(i.e., AOD), a good correlation is found.
For every 10,000 molecules of 02 utilized, abou t one molecule of N20 is produced.
Since there is about one unit of N03 produced for each 10 units of
02 consumed, this means that on the order of 1 N atom in 1000 isconverted to N20 rather than to N03 during respiration.
This N20
appears to be nearly inert in sea water and accumulates until thewater contacts the atmosphere at which point the N20 is lost.
Once in the atmosphere the N 2 0 is gradually converted to N2. Thismechanism alone would destroy all the N03 in the sea in about
1,000,000 years (1000 mixing cycles at 1000 years per cycle).
Not all the N20 produced in the sea escapes to the atmosphere.
At a few places in the sea the oxygen content is drawn
all the way to zero (see figure 3-17).
In the northern part of
the western Indian Ocean there is a layer almost one kilometer
thick with near zero oxygen content.
In this layer nitrate isused as an oxidizing agent.
This can be shown from the very low
NO 3 to P0 4 ratios or by comparing the observed NO 3 content wi ththat predicted by multiplying the phosphate content by 15 (seefigure 3-16). A similar situation is found in the eastern Pacific
Ocean in the 100 to 20 0 N latitude belt (see figure 3-16). In thisregion N20 measurements have been made.
As shown in figure 3-18
the N20 concentration continues to rise with decreasing 02 concentration until O2 levels very close to zero are achieved.
Then,along with the nitrate content, the N20 content starts to drop.
Waters have been found which are nearly free of N20 as well as ofoxygen.
In addition to learning about a sink for N20 within the sea,
we have also found evidence for the destruction of N03 within thesea.
Al though a bit off topic for this chapter (i. e., gases inthe sea), let us explore this evidence for ni trate consumption a
bit further.
It is not only in waters free of 02 that we haveevidence for N0 3 - consumption.
The deep waters of the Bering
Sea which are trapped behind the Aleu tian arc also show evidence
for such depletion.
They are ventilated by deep Pacific wate~s
which enter the basin through a deep gap in the arc near its western end.
The chemistry of wate~s of the same salinity inside andoutside the arc are compared in table 3-6.
It can be seen thatthe Bering Sea waters are, as might be expected, lower in dissolved oxygen and higher in phosphate than their open Pacific counterparts. The ratio of 02 deficiency to P0 4 excess is more or lessas would be expected.
Nitrate, however, is slightly lower inBering Sea ·waters than in the open Pacific. Thus, denitrificationmust be occurring wi thin the deep Bering Sea.
This is not to be
expected, as the oxygen contents are everywhere in the Bering Sea
well above zero.
One explanation for the N03 depletion might bethat the fecal pellets falling into the deep waters develop anaerrobic cores and that these cores then become the sites of nitrateutilization.
In any case, we can add another item to our growing
list of sites in the ocean where N03 is being destroyed.
Whatever p~ocess is operative in the Bering Sea may well be
at work in parts of the open Pacific.
In figure 3-13 it can beseen that the highest phosphate deep waters (i. e., those in the
far northeast of the deep Pacific) show ni trate values falling142

02 (fLmol/kg)

N20 (nmol/kg)
5

10

15

20

140

( 80

220

260

300

E

-.....
~

:J:

a.

w
o

NORTH ATLANTIC

39°N - 62°W
JULY 1972

~8

z

en
en

IJJ

u

~ 00

80
160
A a U(fJ-moilkg),

5.0"---~--"""'--"""--"""----'

Figure 3-15. Plots of dissolved oxygen and nitrous oxide concentration as a function of depth in the northwestern Atlantic Ocean(note the change in the depth scale below 1000 meters).
As shown
by the insert, when the excess N20 content of the samples (i. e. ,measured N20 minus saturation N20 content) is plotted against AOD(saturation oxygen content minus measured oxygen content), an excellent correlation is found.
One molecule of N20 is produced forevery 10, 000 molecules of
2 utilized.
While the ratio of N20
production to 02 utilization varies somewhat from place to placein the sea, except for regions where O 2 is totally depleted, N 20
nuilds up as 02 is used. These results were obtained by Yoshinari
at Dalhousie University in Halifax, Canada (419).

°

143

N0 3 (p.mol/kg)
010',...-.....50 O·°ft_
J

10

,.
E

~

• WESTERNIINDIANoto- 20 N-65°E
•
I

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~

o

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OW

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,

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o

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aBS

\~02

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1.5

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50

PACIFIC
12°N-124°W

~O.5~
- .,

•

I

40

L __ . , -....

EASTERN•I:

0.5 :-

-

30

i(

I

.x:

20
i

50

100

150

02 (jLmol/kg)

200

250

1.5
0

,

•,

50

100
150
200
02 (jLmol/kg)

250

Figure 3-16. Plots of the dissolved oxygen and of the nitrate concentration versus
depth in the northern part of the western Indian Ocean and in the eastern part ofthe equatorial Pacific.
Both areas have thermoclines in which the dissolved oxygen
content goes to zero.
In these waters nitrate is utilized by bacteria as an oxidizing agent leading to nitrate deficiencies in the water column (see the shaded area
between the measured and the expec ted ni t ra te concentra tion curves).
The expec ted
nitrate concentration is calculated by multiplying the measured P0 4 concentration by
15. These results were obtained as part of the GEOSECS program (425, 426).

Figure 3-11. Map showing those regions of the ocean where dissolved oxygen goes to less than 20 ~moles/kg in the main thermocline~
This diagram was constructed by Werner Deuser at the Woods
Hole Oceanographic Institution.

Table 3-6. Evidence for nitrate utilization in the deep Bering Seat

Salinity
0/ 00

OPEN PACIFIC
N0 3
P04
llffi/kg ~m/kg ~m/kg°2

34.625*91100
34.63611034.647124
34.65834.669** 133

2.71
2.69
2.62
2.57
2.53

41. 1
40.5
40.0
39· 3
38.7

°2

~m/kg

BERING SEA
P0 4
NO 3

72
79
87
96
109

~m/kg

~m/kg

2.90
2.89
2.86
2.81
2.75

40.6
40.1
39. 3
38.8
38. 5

BERING - PACIFIC
I::. PO 4
I::. NO 3
1::.°2~m/kg llm/kg llm/kg
-19
-21
-23
-28
-24

0.19
0.20
0.24
0.24
0.22

tBased on the results obtained at GEOSECS stn. 219 (425).
*Depth 2500 meters.
**Depth 3700 meters.

145

-0.5
-0.4
-0.7
-0.5
-0.2

50

EASTERN PAC I FIC
40.......-.

OX Ie

Cl

WATER

~

-'"0

30

E

:L

~

0

20

C\J

Z

10

SUBOXI C

WATER
50

100

150

200

250

02 (fLmol/kg)

Figure 3-18. Generalized diagram showing the trend of nitrous oxide versus' dissolved oxygen in waters from the eastern Pacific(same area as that discussed in figure 3-16).
The usual trend ofincreasing N2 0 concentration with decreasing oxygen content isbroken in waters of very low oxygen content.
Here the N20 is consumed!
These results were obtained by Cohen and Gordon of Oregon
State University (420).

146

well off the trend found for waters with lower phosphate contents.
Again N03- reduction appears to be occurring!
Before leaving the subject of the gases of nitrogen in thesea, we should mention that, whereas the destruction of N03 within
the sea should produce increases in the ni trogen gas content of
sea water, it must be remembered that deep sea water contains
about 600 micromoles of N2 per kg.
The greatest N0 3 deficiency so
far measured in sea water is about 10 ~mole/kg. The corresponding
N2 increase would be only two percent - well within the range ofsupersaturation effects.
EXC_ESS HELIUM

Two gases are being generated within the earth, 40 Ar from theradioactive decay of the long-lived isotope, 40K~ and 4He from the
•
•
L38
235
232
radloacti ve decay of the long-Ii ved lsotopes,
U,
U,
Th
and their respective daughter products.
From our knowledge ofprocesses occurring in the upper atmosphere it is possible to showthat the average 4He atoms are able to "evaporate" from the atmosphere to outer space on a time scale of about one million years.
By contrast the escape time for an 40 Ar atom greatly exceeds theage of the earth. Hence, while virtually all the 40 Ar atoms whichhave reached the atmosphere from the earth's interior are still
there, only those He atoms which leaked from the interior during
the last fe~ million years are present. From our knowledge of theamount of 4 Ar in the air and of the amount of the element potassium in the earth, it is possible to show that a fair fraction (50%) of the atoms of the radiogenic argon produced wi thin theearth have managed to reach the earth's surface.
Based on thisobservation, it seems likely that 4He atoms are outgassing fromthe earth's interior at close to the rate at which they are being~roduced in the earth.
Indeed, the calculated evaporation rate of
He atoms from the to~ of the atmosphere turns out to be similar
to the rate at which He is being produced wi thin the earth.
As
70% of the earth's surface is covered by ocean water, a fair fraction of these 4He atoms must pass through this reservoir on their
way to the atmosphere.
Since the deep ocean waters are replacedon a 1000-year time scale, it was possible to show that this flux
should produce a measurable excess (over saturation) of 4He in the
deep sea. Suess and Wanke (265) were the first to attempt to findthese excesses.
Their resul ts placed only an upper limi t of six
percent.
Subsequent measurements of the absolute helium concentration in deep sea water samples made by Bieri, Koide and Goldberg of Scripps Insti tution of Oceanography (266,267) suggested
that considerable excesses of 4He (5 to 20%) were to be found.
Harmon C~raig of Scripps Institution of Oceanography and Brian
Clarke of McMaster University set out to check the validity ofthese measurements by determining the 3He to 4He ratio rather than
the absolute concentration of helium in deep sea water samples.
The logic was that unlike 4He , 3 He is not beirw produced in theearth's interior.
Rather the trace amount of He found in atmospheric helium was thought at that time to be produced by thespallation of the nucleii of atmospheric gases by cosmic rays.
If
excess 4He of earth's interior origin was indeed to be found in

147

,

x

0

\
\
\

x, ,

,

......

---

J- . . ,,.....,.

E~......"

28°N-122°W

31°S-177°W

2

/
/

I

/

r-

/
/

a..

/

W 3
0

<,'"

/

\

~

4

\

~

~9

1.0

1.1

1.2

3He

3He

4

4He

He

OBSERVED

1.3

1.4

EQU I LBRIUM

Figure 3-19. Devia tions of the 3 He / 4 He ratio from the value for
equilibrium wi th the atmosphere versus water depth at two locations in the Pacific Ocean. These resul ts were obtained by Brian
C.larke at McMas ter Uni ve rsi ty, Hamil ton, Canada on samples collected by Harmon Craig of the Scripps Institution of Oceanography(269,270) .

148the deep sea, then the 3He /4 He ratio in deep sea water should beless than that in atmospheric samples.
Craig collected new deepwater samples and Clarke and his associates analyzed them on a
newly constructed mass spectrometer capable of highly sensitive
and accurate helium-isotope ratio measurements (269) •To their
great surprise they found that rather than having lower 3He /4 He
ratios than atmospheric samples, these samples had higher ratios!
To their glee, they immediately realized that they had made amajor discovery.
3He built into the planet at its beginnings was
still seeping out into the atmosphere!
As their measuremerts 4on
samples from the deep Pacific showed up to 20% higher
He/ Heratios than that for atmospheric helium and since their measurement errors were less than 1%, there was no question that they
were right.
Indeed, subsequent measurements have confirmed thisfinding.
The depth profiles of the 3 He anomalies found by Craig and
Clarke in the Pacific Ocean revealed another bonanza. As shown infigure 3-19, they found that the maximum in this anomaly was notat the bottom of the ocean but rather at about 2 kilometers depth.
This suggested that the excess 3He was coming from the crests ofthe mid-ocean ridges!
Measurements of helium in the recently discovered hydrothermal vents on these ridges reveal very large excesses of both 4He and 3He. The 3He/ 4 He ratio in these hot waters
is eight times higher than that in atmospheric helium.
Since themaximum 3He anomaly shown in figure 3-19 is about 25%, the maximum
4 He anomaly would be expected to be only 3%.
As we shall see inchapter 7, the deep sea 3He anomalies have great importance forstudies of the mixing processes occurring in the deep sea.
Subsequent to the work on 3He anomalies in the deep sea by
Craig and Clarke, a graduate student named Jenkins working in
Clarke's laboratory found another type of 3He anomaly in the sea(271,274).
3He has, over the last few decades, been produced in
the upper ocean at significant rates by the decay of the tri tium
produced during nuclear weapons tests.
This 3He is largely confined to the upper kilometer of the ocean because the tritium produced by nuclear testing has not penetrated in significant amounts
to depths greater than this.
The distribution of this "anthropogenic" 3He allows important constraints to be placed on the mannerin which the main thermocline of the ocean is ventilated. We will
have more to say about this in chapter 8.

THE CARBON DIOXIDE CONTENT OF SURFACE OCEAN WATER
Prior to man's interference wi th the carbon cycle throughdeforestation and the burning of fossil fuels, the C02 content of
the surface ocean and atmosphere was presumably at steady state(i.e., there was no net transfer of C02 gas between air and sea).
We say "steady state" rather than "equilibrium" because sizablevariations in the CO 2 partial pressure in surface ocean water
exist from place to place and from season to season. By contrast
the atmosphere has a more nearly uniform C02 content.
The variations of CD2 partial pressure in the surface ocean
stem from acomplex interplay between the seasonal temperature
cycles, mixing dynamics of the upper ocean, and the biological149

cycles of the sea.
C02 transfer between the air and sea tends toeliminate these differences, but, as the time constant for theachievement of equilibrium between air and surface sea is about 10
times longer for C02 than for 02 (i.e., about 1 year rather thanabout 1 month), this process is not altogether effective.
Onemight say that, although at any given locale the surface water
tries to come to equilibrium with the air, it never quite catches
up with the processes tending to disrupt this equilibrium.
In order to understand this chase, we must first learn what
controls the CO 2 gas content of surface water.
We have alreadyseen that most of the carbon in sea water is in the form of HC03and C03= ions and that less than 1% is in the form of C02 gas.
Throughout the ocean, the following chemical reaction is continually taking place:

3-18
One water molecule, one dissolved C02 gas molecule and one
carbonate ion react to form two bicarbona te ions.
This reaction
proceeds rapidly in sea water andchemical equilibrium ismaintained.
Thus the concentrations of these ingredients obey thefollowing rule:

3-19
where K' is the equilibrium constant for the reaction. *Since
the concentration of water itself does not change measurably as
the result of its participation in the reaction,** the C02
content of any sea water sample is proportional to the square of
the HC03- ion concentration divided by the C03= ion concentration.
We have already seen in chapter 2 that at any place in theocean the HC03- ion concentration and the CO 3= ion concentrations are fixed by the alkalinity, Alk, and the total dissolved
carbon content LC02 of the water.
Using the approximations
adopted in chapter 2 we obtain:(2[LC02] - [Alk]f[Alk] - [LC02]

3-20

The pressu"re of CO 2 gas in air equilibrated with this water willbe:
3-21
*As shown in table 3-7 and figure 3-20, this constant is related
to :he s~ called first and second dissociation constants for carbonlc a~l~. These constants are functions of temperature, pressure
and sallnlty.
In table 3-7 [C02] IS given as the product ofa.peo 2·**Sea water contains 54 moles/kg of H20 molecules and only 2xlO- 3
moles/kg of the various carbon species.

150

Table 3-7. Equilibrium constants forreactions involving thespecies C02, HC03-, C03= and H+. The reactions are as follows.t
C02air ~ C02 seaa = [C02 J/PC0 2
CO 2 + H 20°

H+ + HC 3-¢

aH+ [ C0 3=J[ HC 03-J
C02 + H20 + C03=
K'~

2HC03aK 1 '[HC0 3 -J 2

=--

K2 '

Pc

°

2

[ C 03=J

For sea water of salinity 35. 0 /00 at 1 atmosphere pressure theconstants are as follows:
T

°c
o
5

10
15
20
25
30

a*

aK 1 '/K 2'

moles/kg atm

6.29xl0- 22
5.21xl04.39xl0- 2
3.74xl0- 2
3.24xl0- 2
2.84xl0- 2
2.52xl0- 2

moles/kg atm

6.33xl0- 77
7.17xl07.98xl0- 77
8.73xl09.41xl0- 7
10.00xl0- 7
10.47xl0- 7

103.53xl0- 10
4.04xl04.72xl0- 10
5.58xl0- 10
6.59xl0- 10
7.69xl0- 10
8.77xl0- 10

113
92
74
59
46
37
30

*Based on measurements by Ray Weiss of the Scripps Institution of
Oceanography (403).
**Based on measurements by Merbach and his coworkers at Oregon
State University (400).
taH+ is the activity of hydrogen ion.

151

7~

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1

1

i

Jdll

r

AT I ATM

13

-E 5
.....
~

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.........

0

-

0

.§
tj

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,.....---

0'

8-~

4

.......

i

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7

9

0

1

0

6'
0

I

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I

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-15

..-

10

0

0

0

---

--- --- 9

-N
~

-

-N
~

~

4

5

CJ1

]6

AT 2°C

I
I

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12

8

10

~

N

8

3r- / /
2'

o

,

I
(1

,
10

,

~

17

,

16
30

,
20

TEMPERATURE (OC)

7

3

6'

o

,

,
200

,

,
400
PRESSURE (atm)

,

'3
600

3-20. Equilibrium constants (at 35%0 salinity) for the carbonate system
in sea water.
The left hand panel shows the values of the constants for surface
ocean water as a function of temperature.
The right hand panel shows the pressure
effect on the constants for cold sea water (i.e. at 2°C).
The dissociation constants (K l' and K2 ') were determined by Merbach and his coworkers at Oregon State
University (400).
The solubility of CO 2 gas was determined by Weiss at Scripps Institution of Oceanography (403).

Figurewhere a is the solubility of C02 in sea water at the temperatureof interest (see table 3-7).
Combining with the previous equation, we obtain:

PC0 2

1

= aK'

[

(2[LC02] - [Alk])2 ][Alk] - [LC02]

3-22

Since both a and K' depend on temperature and salini ty, so does
the equilibrium C02 partial pressure.
'l'he temperature dependence of the equilibrium constant product raises a problem.
Were it true that all surface water had
lost its phosphate and nitrate to plants and that this loss
brought the alkalinity and total dissolved C02 content to uniform
values, then the [HC03-J2/[C03=] ratio would be nearly constantthroughout the surface ocean.
However, even in this idealized
ocean the re would be a problem.
Since aK' varies wi th temperature, the equilibrium C02 pressure for any given area of the ocean
would depend on the temperature of the water.
As the equilibrium
CO 2 pressure rises 4 percent per degree-centigrade increase intemperature, the equilibrium C02 pressure for tropical water would
be about 2.5 times as great for polar water!
This would tend tocreate a poleward decrease in atmospheric C02 content.
However,
the atmosphere mixes so rapidly that no such gradient would be
generated.
If, then, the C02 partial pressure for polar surface sea
water is lower and the CO 2 partial pressure of tropical wa ter ishigher than the steady state CO 2 pressure for the well-mixed atmosphere, CO 2 must be flowing out of tropical surface water and into
polar surface water.
This flow would then disrupt the geographic
uniformi ty of LCO 2 in our idealized surface ocean.
We can thus
consider two extreme cases, one where this transfer is ineffective and the CO 2 partial pressure in tropical surface waters is
far higher than that in polar surface waters; and the other where
the transfer goes to completion, alleviating the C02 partial pressure difference between these two regions of the ocean.
For thelatter case the LC.o2 concentration in surface would increase toward the poles in accord with the following equation (obtained bysolving equation 3-22 for LC02):[l:C02J ;~ [4[AJ-aK' Pe0 2+/( aJ(' PC02 )-4[AJ) 2- 1 6[AJ( [AJ-aK' PC0 2 ) J

3-23

In this equation A denotes Alk. These two extreme scenarios areshown diagrammatically in figure 3-21.
Where in the range between these extremes would we expect thesituation in the real ocean to lie?
To venture this judgment wemust know the response time for the adjustment of the EC02 content
and C02 partial pressure in the surface ocean.
We learned earlier in this chapter that the time required for an ordinary gas toachieve equilibrium between the atmosphere and the surface ocean
mixed layer is a few weeks.
If this time were also to apply to
CO 2, then gas exchange would be able to keep pace with the153

............;J

p~~~.~.~.~~. ~~~

pi

LC02

t

t

---;------

C\.I

o

u

C\.I

o

U

~

Cl.

PC02 AI RSLUGG ISH EXCHANGE

60

o

S

LATITUDEe.

......

I-. ..

a. .. . ... ...

~t..... ...

.. ..... .... .. ....... .... .. ..

PC0 2 AIR
]

LC02

C\.I

o

-

-----~---

u

60
N

7

Cl.

.-

..
....

.-

t

_ _ _ 0C\.I
U
V'J

PC02 SURF H2 O
RAPID EXCHANGE
I

Figure

60

0

60

S

LATITUDE

N

3-21. Extreme

scenarios

for

the

latitudinalvariationofPea z and Lea 2 in the surface ocean (assuming constant alkalini ty) •154

processes tending to drive the C02 pressure in a given region of
the surface ocean away from the atmospheric value. We should havea surface ocean which had a nearly uniform C02 partial pressure
and a >::C0 2 concentration increasing from equator to poles.
The
equilibration time for CO 2 is not the same as that for othergases. The reason is that in order to change the partial pressureof C02 in surface water, not only the CO 2 content of the watermust change, but also the HC03- and C03= ion contents.
Since
the EC02 content of surface sea water is about 150 times the C02
gas content, one might be tempted, then, to say that it would require 150 times longer to adjust the C02 partial pressure in surface water than, let r s say, the 02 partial pressure.
This would
also be wrong.
Actually, it takes about 10 times longer to adjust
the C02 partial pressure than that of ordinary gases.
To understand why this is the case is not easy.
In fact, it
involves one of the more complex aspects of marine chemistry.
We
will delay our consideration of the precise explanation until
chapter 10 where we deal with the uptake of fossil fuel C02 by theocean. At this point it will suffice to say that it is not the C02
or >::C02 contents of sea water that set the equilibration time but
rather the C03= ion content of the sea water.
As the concentration of carbonate ion in surface water averages about one tenththe >::C02 content, the multiplier is closer to 15 than to 1501
It is the CO 3= ion content which is important because themain reason C02 enters and leaves sea water with, lets say, changing temperature is through the reaction:

3-24
This reaction between CO 2 and C03= leads to much larger changes
in EC02 content than does the change with temperature of the solubility of C02 gas.
In any case, the response time of the PC0 2 and EC02 in surface ocean water averages about one year.
An example of the response of the pC02 and EC0 2 content of surface water to a sudden
temperature change is shown in figure 3-22.
Clearly, this time is
too long to allow the ocean to keep exact pace with seasonalchanges. Further, as surface waters in regions of rapid upwellingand in regions of deep water formation are not in contact with theatmosphere for a time as long as one year, these waters would not
be expected to achieve equilibrium with the atmosphere.
In figure 3-23 are shown the temperature, salinity, ECD2 andalkalinity. along a north-south traverse in the western Pacific.
As can be seen, the ECD2 content of the surface waters does increase from temperate to polar regions.
However, contrary to theassumptions regarding our idealized surface ocean, the alkalinityof the water increases as well.
Thus it is not immediately clear
how much of this increase is due to temperature and how much is
associated with the higher alkalinity of polar waters (as for nitrate and silicate, the alkalinity of Antarctic surface waters ismore akin to the alkalinity of deep water than to that of warm
surface water).
Also seen in figure 3-23 is a curious peak in
EC0 2 content at the equator.
In this case there is no matchingalkalinity anomaly.

155

~

U
......-

20

0

~

~

19

W

I-

0.0

0.5

1.0

0. 0

0.5

1.0

1.5

10.00

0.0

0.5

1.0

1.5

10.00

320
~

E
c

.p-

318
316
314

to
I

0

.........,.
(\J

0

312

c..

310

U

4%

308
.........

2.008
rt')
E 2.006
"0
2.004

E

......-

(\J

2.002

0

2.000

W

1.998

u

TI ME (years)
Figure 3-22. Response of a 50 meter deep thick surface mixed layer
initially at equilibrium wi th the air to a sudden cooling from
20 0 e to 19° C.
Its PC0 2 drops by 4 percent (12 x 10- 6 atm).
Theinvasion of atmospheric C02 then causes the LCO 2 content to rise
and the Pea 2 to move back toward the atmospheric value (320 x
10- 6 atm).
The half-time for the reestablishment of air-water
equilibrium is about 0.4 years.

156

2300

o '\.

2200

"
\

•\

-

2380

/~2360 doo<V\~ 2100
o

\

*C\I

•
~
,
~ ,

r.C02*/

1950

2340

I

\0

--2000
8
H

___

:

. \

~2050

::L

II

\

E

~

I

35.00 %0\

0'.

2320 Z

~

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/

/0'0

P

0,

I

\

2300 ~//0\0 /' .....0 _ _ .; '/0

>-

l-

I
"
I
0 ....00....0 -uJ -*NORMALIZED TO
A SALINITY OF

\0

,-r
'l.

2280

1900
30 r----.---r--....--.--..,...---.....-..~ .....---r--....-.....---.360"""-""

24

w

0::::

::J

\

18

I0::
W

WESTERN PACIFICSURFACE WATE RS

Q..
~

w

l-

I

40

S

Figure34

TEMPERATURE

<t

20

0

LATITUDE

3-23. Plots

,

20

N

of alkalini ty (normalized to a salini ty of(normalized
to asalini ty of
35 0 /00) ,tempera ture and salini ty for surface waters in the wes tern Pacific
Ocean as a function of latitude.
Strong equatorial anomalies areseen in ~C02, temperature and salinity but not in alkalinity.
Based on results obtained during the GEOSECS program (406).

35 0 /00) ,

~CO 2

157

In figure 3-24 are shown the CO 2 partial pressures measured
along the same traverse.
As can be seen, there is remarkably
little variation from the temperate to the polar regions.
Apparently these waters have remained in their respecti ve zones longenough to achieve more or less the ambient C02 partial pressure of
the atmosphere.
We say more or less because, even after steadys tate has been achieved, seasonal temperature changes will dri ve
the CO 2 partial pressure back and forth across the atmosphericvalue.
The startling feature of this diagram is the strong peakin C02 partial pressure at the equator.
It nicely matches thepeaks in LC02, temperature and salinity seen in figure 3-23. Hereis an example of surface waters which have not had the time toachieve equilibrium wi th the atmosphere.
As we shall see, thisfeature is a product of the strong upwelling along the equator in
the Pacific Ocean.
Also shown in figure 3-24 are the C03= ion content of surf ace water and the values of a parameter termed the Revelle factor. * As can be seen, one has roughly the inverse shape of theother.
The Revelle factor is the number by which the ratio of themixed layer thickness to the piston veloci ty for CO 2 gas must be
multiplied in order to get the actual LC02 equilibration time.
This number ranges from 14 for the coldest waters to 8 for thewarmest waters in the surface ocean.
Its ocean-wide average of 10is the factor we referred to above as the difference between theequilibration times for C02 and other gases.

ORIGIN OF THE EQUATORIAL PACIFIC C02 ANOMALY
The surface waters of the equatorial Pacific profile provide
an excellent example of the interaction between ocean dynamics,
meteorology, the activity of organisms and gas exchange.
As seenin figure 3-23 surface water in the equatorial Pacific also has alower temperature and higher salinity than the adjacent waters aswell as anomalies in C02 pressure and the LC02 concentration.
Further, as shown in figure 2-2, sediments of very high opal content are found immediately beneath this belt.
All of these features can be explained by upwelling.
The oblique convergence of the trade winds onto the equatorial zone leads physical oceanographers to predict a divergence ofthe surface waters along the equator.
Filling the "gap" left bythis divergence are waters rising from beneath the surface. Thesewaters are colder than the air over the equatorial zone (leading
to the low temperature anomaly); they carry higher nutrient constituents ~eading to high plant productivity (especially for diatoms) and high opal accumulation rates, and they carry water highin LC02 content leading to the anomalies in LC02 and in PC02·
While the depth from which the upwelling waters are drawn is
not known, waters from any depth below about one hundred meters*Named after Roger Revelle, who while serving as the Director of
the Scripps Institution of Oceanography, pointed out the importance of this parameter in the calculation of fossil fuel C02 uptake by the ocean.

158

500,........,.--r--.,-.....,-..,..-..,....---,r-----ro-..,..-......- .__- - - UNDERWAY MEASUREMENTS OF Pe02 IN AI R
- - UNDERWAY MEASUREMENTS OF Pe02 IN SURFACE WATER•
PC02 CALCULATED FROM ~C02 AND ALKALINITY
M~ASUREMENTS ON DISCREET WATER SAMPLES

-

o
0

~
1

-C\J

o

~ 300

••••

•

•

• •

WESTERN PAC I Fie.....__.-_.. 14~-r-...,..--r--.,....-~---.--r"-.,..-.._-

300

13

a::

0
12 I-

Cl

.:s:. 250u........

o

Li:

~ 200

-

II

uo 150

10 W

w
-!
-l

1 Irt')

>
w
9 a::

100

50 "'--.Io......._ ___~ ~ - . I o - ~ ......~ ~ -........-.~-- 8

~-

S

N

Figure 3-24. Plots of the C02 partial pressure, of the eo 3= ionconcentration and of the Revelle factor in surface waters from thewestern Pacific as a function of latitude.
Two measures of thesurface water Peo 2 are given.
The solid dots are PCO 2 values
calculated. from tltration measurements of 2:C02 and alkalinity.
The solid line gives the smoothed trend based on continuous measurements of the CO 2 pressure in air equilibrated on shipboard withsurface sea water.
The latter are more precise and accurate than
the former.
The Revelle factor is an exact measure of the ratioof time required for the CO 2 pressure in surface water to equilibrate with the overlying air to the time required in sea watercontaining no bicarbonate or carbonate ion.
The Revelle factor
shows an inverse correlation to the carbonate ion concentration.
The Revelle factor is half again as large in the very cold Antarctic surface waters as it is in low latitude surface waters.
Basedon results obtained as part of the GEOSECS program (406).

159

LATITUDE

S
10

0

0

N
10

20
4005001000"",
ISO

-

0"

-E
I

0°

I0..

w 800

0

/'~

1000
~Q

1200

°0

POTENTIAL PC02

"---

1400

Figure 3-25. Potential C02 pressure as a function of depth in theequatorial Pacific.
Potential CO 2 pressure is the CO 2 pressure asample would achieve if it were isochemically depressurized (to 1atm) and warmed (to 25°C).
Based on results obtained by Dave Basof the Scripps Institution of Oceanography as part of the HawaiiTahiti Shuttle Experiment (433).

160

could do the job.
As shown in figure 3-25, all these waters havepotential C02 partial pressures* well above those seen in thesurface waters of the equatorial zone.
These very high pressures
are not seen at the surface
because after their arri val in themixed layer plants remove carbon to form their soft tissues andCO 2 is lost to the air.
Both processes lead to a reduction in
LC02 and in PC0 2 of the water.
Thus part but not all of thePC0 2 and LC02 anomalies are eliminated during the residence of
this water at the surface.
As shown in figure 3-23, there is noequatorial anomaly in alkalinity.
The plants growing in equatorial waters are mainly diatoms and the herbivores grazing on theseplants do not live in calcite houses.
It is possible to separate the role of plant growth and of
CO 2 evasion in reducing the excess carbon content of the upwelling
water through plots of EC02 concentration (normalized to a salinity of 35.00 0 /00)
against nitrate concentration for samples
taken at various water depths and locations in the equatorialzone. As can be seen in figure 3-26, the points for samples takenin the upper thermocline follow a trend consistent with the biochemical ratio of carbon to nitrate in marine organisms.
Pointsfor surface and near surface waters, however, deviate from thistrend.
They have lower EC02 contents than would be expected fortheir nitrate contents.
This missing LC02 has presumably been
lost as C02 gas to the overlying air.
We will pick up this thread again when we consider the
problem of harnessing the information carried by ocean tracers toyield rates of thermocline ventilation (chapter 8).
For now wewill only point out that the upwelling process must be sufficiently rapid that the waters of the surface equatorial ocean are
replaced on a time scale short enough to prevent either plants or
gas evasion from completing the job of bringing the CO 2 pressureback to the atmospheric value. While we do not have a good handle
on how long the plants living in equatorial surface water take tocomplete their task, we do know how long it would take for gas
evasion to eliminate the C02 over pressure.
This could be accomplished in about one year.
Thus the upwelling process must
operate at the rate of several tens of meters per year in order tomaintain a C02 anomaly in the equatorial mixed layer!

SUMMARY
We have seen that the gases dissolved in the sea add an important dimension to marine chemistry.
From measurements carriedout in the laboratory their solubilities in the sea water are
known and from measurements of radioactive gases in the air andsea their rate of transfer across the air-sea interface is also
known.
The latter is sufficiently rapid that the concentration of
all gases (except CO 2 ) in waters descending to the deep sea can beassumed to be qUite close to equilibrium with the atmosphere.
Thefact that this is the case allows us to put to effecti ve use*The C02 partial pressure that would be achieved if the water werebrought to the surface and isochemically warmed to the ambient
equatorial ocean temperature.

161

2300

..

2300

Ite

0'

~

.......

2200

o

E

3- 2100*ON 2000

2100
2000

U

t!

1900 •~o_ _4 -_ _10~_ _-J,.._ _-...J,_....j.

20

30

*

- ' -_ _-.l-1900 l -_ _--J-

40

10

0

(N0 3 ) (p.mollkg)

20

30-.J._....j.

40(N0 3 )*{l-LmOl/kg)

0'~.......
o2200

E

3- 2100
*N
o
U

..

2000

~

1900I.-.._ _-I.-_ _- ' -_ _--l-_ _---l-_-I-

o

20

10

30

40

(N03)* (p.motlkg)

..

2300

• e

0'

.:.::

.......

.

r-----t-----t-----t------r---r

: _

2300

e

2200

"0
E

:L

2100

*N

2000

0

u

•

~
1900
0

10

20

30

40

1900 I - -_ _-I.-_ _-I--o--<_--l-_ _--'-~
40
10
20
30
o

(N0 3 )* (p.mollkg)(N03)* (/J-mol/kg)* NORMALIZED TO A CONSTANT SALINITY
Figure 3-26. Plots of EC.02 (normalized to a constant salinity) against nitrate for samples of the surface and thermocline water infive latitude belts in the equatorial Pacific (at longitude 155°W).
In each case a straight line is fit to the points with intermediate phosphate content.
As in each case this line has a slope
close to the Redfield ratio (-6.8), it is assumed that these samples have mainly been influenced by the photosynthesis-respiration
cycle.
For shallower waters (i.e., those with less P04) a deviation toward lower ECD2 contents is seen.
This deviation is presumably due to CD2 loss to the atmosphere. For deeper waters (i.e., those with high phosphate content) a deviation is also seen.
This deviation can be shown to be the result of calcite dissolution as its ons et corresponds to the depth where the alkalini ty
begins to rise.
Based on resul ts obtained by Dave Bos of the
Scripps Institution of Oceanography as part of the Hawaii-Tahiti
Shuttle Experiment (433).
The track followed during this expedition is shown on the foldout map.

162

the distributions of 02, N20, and 3 He within the sea.
In laterchapters we will show that the distributions of two tracer gases
not yet discussed in this chapter (39 Ar and 85 Kr ) will eventually
allow a great improvement in our knowledge of the modes and rates
of ventilation of the thermocline and deep sea.
Finally, thesea's most interesting consti tuent, carbon, is partly in gaseousform.
Our knowledge of the gas exchange process will help us tosort out its movements through the ocean-atmosphere system.
Itwill also allow us to make more effective use of the distributionof the most important tracer in the sea - namely, radiocarbon.
PROBLEMS

3-1

A rapidly stirred tank of gas-free water at 2°C is exposed tothe air.
Its depth is 3 meters.
After an hour, the 02 concentration is found to be 10 percent that expected for equilibrium with the atmosophere.
What is the stagnant boundary layer thickness for the water in this tank?

3-2

A sea water sample is quickly heated from 15°C to 25°C.
If
it initially had a Pea
equal to that in the atmosphere(3.2xl0- 4 atm), what woul~ the partial pressure of CO 2 gas beonce it was heated?
The alkalini ty of the water is 2.2 x
10- 3 eq/kg.
(Assume that no biological activity or gas
exchange occurs during the heating.)

3-3

Mono Lake in California has a very high I:CO 2 concentration(0.2 moles/Ii ter) .
The lake has a mean depth of 30 meters.
Its surface temperature averages 18°C.
The 14 C/ C ratio in
its dissolved carbon is uniform with depth and is 30 percent
lower than that expected for equilibrium with the overlying
air.
If the 14 C deficiency is at steady state, what must thestagnant film thickness for this lake be?
The input of new
carbon from the rivers supplying the lake is too small to besignificant.
(Hint:As in the ocean, the loss of 14C byradiodecay wi thin the lake InUS t be balanced by the net influx of l~C from the air.)

3-4

A tank held at 24°C has a mean depth of 3 meters and a stagnant film thickness of 40 microns (as measured via argon
evasion experiments).
If this tank of water is charged with
0.5 atm of He and 0.5 atm of Ar and then allowed to standopen in contact wi th the air, what percent of the Ar will
have "been lost when 95 percent of the He has been lost? Whatis the ratio of the "piston velocities" for the two gases?
How would the resul ts differ if the film replacement model
rather than the stagnant film model were used? For the film
replacement model the flux is gi ven by the following equation:
F(Csat - Cwater)

163

where T is the residence time of the film at the surface(i.e., T replaces z as the parameter characterizing theexchange rate).
The dependence on molecular diffusivity
changes from the first power to the one half power.
3-5

Equal portions of two water samples (one at 5°C and one at25°C), at equilibrium with the atmosphere are mixed in such a
way that no gas is lost or gained. The concentration of eachatmospheric gas in this mixture is then compared with that ina water sample at 15°C which has been brought to equilibriumwith the air.
Differences are found.
Why?
Would theanomaly be larger for helium or for xenon? Why?

3-6

Air bubbles are injected into the base of a 2 meter highcolumn.
The bubbles enter a trap and eventually dissolve
into the water.
The injection continues at a constant rateuntil steady state is achieved (i.e. the excess gas supplied
from the bubbles leaves by ordinary gas evasion as fast as
it is being added wi th the bubbles.).
At this point 4 He is
found to be 20% supersaturated.
By how much would 02 besupersaturated if the stagnant model applied?
If the film
replacement model applied. The temperature is 24°C.

SUPERPROBLEM#3

A geochemist, Carlos Chilling, sets out to experimentally
check the difference in the equilibration times for 14 C02 , C02 and
02. To do so he builds a large column which has a cross-sectional
area of one square meter and a height of 11 meters.
He places one
cubic meter of sea water in the bottom of the column. He then installs a fan in the air space above the water designed to agitate
the air-water interface and to keep the air mixed.
He also installs a stirring device to keep the water mixed.
He insulates
the column and installs a temperature control system which maintains the temperature at 24°C.
Chilling injects a touch of
mercuric chloride into the sea water to stop biologic activity.
After three weeks he measures the ~C02 in the water, and thePC0 0f the water and air.
He gets the following results:
2[~C02J = 2,000 x 10- 3 moles/liter

PC02
PC0 2

air
=

water

330 X 10- 6 atm= 330 x 10- 6 atm

He also measures the 222 Rn activity of the air and the water andthe 226 Ra activity in the water.
The results are as follows:
A222 Rn
A222

A226

.
alr

=

0.50 dpm/IOO liter

Rnwater

0.50 dpm/lOO liter

Rawater

5.50 dpm/100 liter

164

From this data he calculates the apparent stagnant film thickness.
What does he get?
Now Chilling is ready for the main experiment.
He replaces
the air with pure nitrogen.
He replaces the nitrogen in the sea
water with pure oxygen.
He then injects some 14 C tagged C02 gas
into the sea water (mostly 12 C02 with enough 14 C02 . atoms present
to swamp out any initially present in the tank).
During these
operations Chilling seals off the air space so that no exchange
can occur.
When he has finished these injections he measures thepartial pressures of O2 and of C02 in the air and water. He finds
the following:
P02water
PC02water=

1.0 atm

350 x 10-6

atm

Po 2 air•
PC0

2

alr

0.0 atm
330 x 10- 6 atm

Chilling pulls the baffle separating the air and water andturns on the fan and stirrer.
He then monitors the 14 C/ C ratio in
the air and in the sea water at frequent time intervals.
He alsomeasures the partial pressure of C02 in the air and in the waterat the same intervals. Finally, he measures the 02 content of the
air and of the wa ter at the same intervals.
Eventualy he findsthat these parameters cease to change.
At this point what shouldbe the PC02 and P02in the air and what should be the ratio of
14C/ C in the water to that just after spiking?
Now Chilling plots the PC0 2 in the air, the P02in the air
and the 14C/ C in the air against time.
For each he finds a
response such that the difference between the initial and steadystate value is reduced exponentially with time (asymptoticallyapproaching steady state). For each property he determines a half
response time.
What values should he find?
In the calculations use following relationship:~PC02/PC02

9.0

l\~C02/~C02

The solubili ties of oxygen and radon are gi ven in tables 3-1 and
3-2.
The diffusivities of 222 Rn , 02 and C0
are given in table
3-3. Assume that the diffusivities of 1 4CO 2 and 12 C02 are
identical.

LAST MINUTE ADDITION (SEPTEMBER 1982)
James Ledwell of Harvard showed, as part of his Ph.D. work,that the gas exchange rates of He, CH 4 and N2 0 in a wind tunnel
follow the square root of diffusi vi ty dependence as predicted by
the film replacement model (721).

165

ekapte/t 4REACTIVE METALS AND THE GREAT PARTICULATE SWEEPTHE CYCLE OF METALS IN THE SEA

INTRODUCTION
In our discussion of chemical cycles within the sea we rather
quickly passed by those elements which the chemis t classifies asreactive metals.
Some of these elements fit into the non-cycling
ca tegory of chapter 1.
They become firmly a ttached to particles
and are removed to the sediment during their first pass throughthe sea.
Others fall into the biolimi ting and biointermediate
categories.
Our focus in this chapter will be on the marine
chemistry of these constituents of sea salt.
The geochemistry of metals in the sea has only recently begun
to be unraveled.
The reason is that these elements are present insuch low concentrations that, prior to this, little success wasachieved in their measurement.
Only after 1975 did reliable resuI ts begin to appear in the Ii tera ture.
These new resul ts reflect a combination of care to avoid contamination during collection, storage and measurement and the use of highly sensi ti ve
analytical methods.
While prior to 1975 the sea water concentrations of these
metals were not known, the concentra tions in marine sediments
were.
These measurements showed that the red clays of the deep
ocean differed from the detri tal rna terial accUffiula ting on continental margins and in embayments of the sea.
Red clays havehigher concentra tions of Fe, Mn, Ni, Co and Cu than do shales(indeed their red color is related to the presence of ironoxide) .
These same elements make up manganese nodules.
Marinescientists have long puzzled over the chemical pathways leading to
the enrichment of these particular elements in deep sea deposits.
To P4t these excesses in context, we will have to review a
bit of the budget of elements in the earth's crust.
Of the igneous rocks found at the sut'face of the earth, the granitic andbasal tic varieties are the mos t abundant.
It is reasonable toassume that they are the dominant sources for sediments. The mostabundant sediment is shale; it constitutes about 80 percent of thetotal.
A comparison of the composi tion of shale wi th that ofaverage basal t and of average grani te is shown in table 4-1.
As
can be seen, a mix of 50 percent grani te and 50 percent basal tyields the concentra tions of mos t elements found in shale.
rrheelements Na, Mg and Ca are exceptions but can be readily explained.
Calcium and magnesium are deficient in shale because
theyresidein limestone(calciumcarbonate)anddolomite

166

Table 4-1. Comparison of the chemical composition of average

igneous rock (equal parts granite and basalt) with the chemicalcompositions of shale and red clay. (Concentrations are inparts per million by weight.) Note the large excesses of Mn,
Ni, Co, Cu and Fe in red clays.

Element

BasaltGraniteAverage
Igneous
Rock

Na
Rb
Cs

26,0008,000
30
118,00042,000170
422,00025,000100
310,00027,000140
515,00025,000
110
6

Mg
Ca
Sr
Ba

46,00076,000
470
330

1,600
5,000
100
840

24,00040,000290
590

15,00022,000300
580

21,00029,000
180
2,300

Fe

86,000
1,500
130
50
90
170
4

14,000
400
5
1
10
4
18

50,000
950
70
25
50
90
11

47,000
850
70
19
45
90
12

65,000
6,700
225
74
250
90
12

K

Mn

Ni
Co
Cu
Cr
Th

167

Shale

Red
Clay

(calcium-magnesium carbonate) as well as in shale.
If we were to
add all Mg and Ca contained by dolomi te and limes tone to theshales, their magnesium and calcium contents would become nearly
equal to the value expec ted from a 50-50 grani te and basal t mixture.
Sodium is in short supply because much of the sodium
released during weathering remains dissolved in the sea.
If wewere to take all the sodium out of sea water and incorporate it
into shale, its sodiwn content would be brought close to the
expec ted value.
When we compare the composi tion of deep sea red clays wi ththat of shale, the agreement is good except for barium and theelements enriched in manganese nodules (i. e., Fe, fvrn, Ni, Co andCu) .
It should be emphasized tha t this enrichment in deep sea
clays relative to shale does not apply to all trace metals.
Chromium and thorium are examples of metals tha t do not show ananomaly.
.
.The rVIn content of a mixture of equal amounts of grani te andbasalt would be about 950 parts per million (ppm).
This is also
the average value found for shales.
By contrast the Mn content of
deep sea sediments averages about 6700 ppn (a factor of 7 higher
than expected).
Ni is, on the average, a factor of 3 higher inconcentration in deep sea sediment than in shale; Co is a factorof 4 higher; Cu is a fac tor of 5 higher.
Iron shows abou t 40
percent enrichment in deep sea sediments over what would be expec ted from the grani te-basal t average or from shale.
Thus thesame processes that caused the enrichment of Mn and Fe in nodules
and in deep sea sediments appear to be at work for other metals.
Our task in this chapter is to learn more about how these andother metals move through the sea.
Our informa tion regarding the pa thways followed by metal scomes from three major sources.
The first of these is the
distribution of radioisotopes produced by radiodecay within thesea.
These isotopes are all members of the decay chains of thelong-lived isotopes 238 U , 235 U and 232 Th (see figure 4-1).
Theisotopes of interest are, of course, those of elements which areadsorption-prone (i. e., thorium, protactinium, lead and polonium) .
The second source of informa tion is the dis tribu tion ofadsorption-prone radionuclides added to the sea surface from theatmosphere.
Plutonium produced during nuclear testing is one
example; the lead produced by the decay of the radon released tothe atmosphere from continental soils is another.
The final
source of .informa tion is the dis tribu tion wi thin the sea of themetallic elements themselves.

PRODUCTS OF URANIUM AND THORIUM DECAY
The marine geochemistry of daughter products of the uranium
and thorium series appears at first glance to be qui te complex.
As shown in table 4-2, the activities of successive daughter products of a given decay series are far from equal (as radioactive
equilibrium would dictate).
Rather large jumps in activity between sUbsequent products are seen..
As shown in table 4-2, thispa t te rn d iffe rs in detail from wa ter ty pe to water type!
Thesource of these differences is as follows.
The isotopes of ab-168

U-238 Series

Element

Th-232 Series

U-235 Series

Neptunium

Uranium

Protactinium

Thorium

Actinium

Radium

U-234

U-238

I

lJ-23S

2.48XI0 5. / YI~

4.47 X 10 9
YIS

Pa-234'1.18--,01107.04 X 10
yrsI

Th-234....

Th-23024.1
days

7.52 X 10 4
yrs

1.40 X 10'0
yrs

I

I

Th-232

Francium

en
("C)

Radon

Astatine

Lead

I

I

Ae-227 ....
21.8
yrs

Th-227
18.7
/days

I

Ra-223

1.62 X 10 3

5.75
ylS

3.66
days

11.4
days

I

I

I

Rn-222

Rn-220

Rn-219

3.82
days

55.6
sec

3.96
sec

I

3.05

I

Po-210Po-2161.64 X 10" 4138........ days

0.15
sec

I

I

1/ sC'c
Bi-214....
19.7

l / minI

1'0-214I

Bi-210......
5.01
/days

Pb-214.....

"b-210....

Pb-206

Pb-212....

26.8

22.3
yrs

slable lead(isolope)

10.6
hrs

min

Thallium

25.5
hrs

Ra-224min

Bismuth

Th-231/

1.91--,yrs4Ra-228.......

Po-21R
Polonium

6.13
/hrs

Th-228

Pa-231
3.25 X 10
i/ yrs

R.1-226
YI~.......

Ae-228.....

I

8

1
64 %

Bi-2 I 2""'"60.6..........minPo-212

1'0-215

3.0X 10- 7./"'sec

1.78XI0- 3

I

I

Pb-20B36

0/01stable lead
I..(isotope)s~cPb-211~36.1
minBi-211
2.15
./min

I

TI-20R ......

TI-207

3.05

4.77

min

min

Pb-207stable lead[.,fisutope)

Figure 4-1. Chart showing the decay chain of the uranium and thoriwn series isotopes
and the half-lives of each isotope.
Alpha decays are shown by the vertical arrows
and beta decays uy the diagonal arrows.

Table 4-2. Summary of the approximate activities of the parent anddaughter isotopes for the three-decay series in various majorwater types of the ocean.

N. Atl.

BottomWater
dpn/100kg

Antarctic
Bottom\~a ter
dpn/100kg

N. Pac.
BottomWater
dpn/lOOkg

240
240
280

240
240
280

-

-

13
>13
8
8

20
>20
10
10

240
240
280
0.15
34
>34
16
16

13

13

13

-

-

-

<0.1
3
0.4

<0.1
O. 4
0.3

Daughter
Produc t
Half-lifeyr

Harm
SurfaceWater
dpn/lOOkg

238U
234 Th234U
230 Th226Ra
222 Rn210Pb
210 poParent0.066
248,000
75,200
1620
0.010
22. 3
0.38

240
230
280
<0.02
7
5
20
10

235 U
231Pa

Parent32,500232 Th228 Ra228Th

Parent5.8
1.9

Isotope

<0.1

-

13
0.05
<0.1
0.4
0.3

Table 4-3. Soluble parent-reactive daughter pairs.
PARENT
Isotope
238 U
228 Ra234 U
235U
226 Ra210 PbHalf-lifeyr
4.47x10 95.75
2.48x10 57.04xl0 8
1. 62x 10 3
22.3

DAUGHTER
Decay
Cons tant
yr- 1
1. 55x 10- 10
0.120
2. 79x 1 0- 6
9. 84xl0- 1 0
4.28xl0- 4
3.10xl0- 2

Isotope
234 Th228Th
230 Th
231 Pa210 pb210 po170

Half-life
yr
0.066
1.91
7.52xl0 4
3.25x10 4
22.3
0.38

Decay
Constant
yr- 1
10.5
0.363
9.22x10- 6
2.13x10- 5
3.10x10- 2
1.82

sorption-prone elements genera ted wi thin the sea are rapidly removed to the sediments.
Isotopes of more soluble elements areproduced in turn within the sediments and leak back into the sea •.The great value of the absorption-prone daughter products tomarine chemistry lies in the fact that their main input to the sea
is via the in situ decay of their immediate radioactive parents.
This production rate can be precisely determined by measuring theconcentra tion of the parent in the wa ter mass of interes t.
Bycomparing the in si tu concentration of the daughter isotope withthat of its paren~e not only can show whether the daughter isbeing removed onto particulate matter, but also at what rate thisremoval occurs.
The parent-daughter pairs of particular interest
are listed in table 4-3.
The principle behind the use of these pairs is quite simple.
If significant removal of tne daughter produc t occurs during its
lifetime, then the radioactive decay rate of the daughter will beless than that of the parent.
The extent of this deficiency is ameasure of the removal rate.
A very simple model of this process can be constructed if it
is assumed that in situ production by the decay of the parent is
the only source Tor the daughter in the wa ter mass of interest.
In this case a "scavenging" probability equivalent to the radiodecay probabili ty can be used to describe the removal process.
Abalance between production by the radiodecay of the parent, scavenging of the daughter product by particles and radiodecay of thedaughter product can be written as follows:
Ap[P]= AD[D] + feD]

4-1

where Ap and AD are the decay cons tants for parent and daught e r, f is the 1f s c a veng i ng " cons tan t *, and [ p] and [D ] are th e
concentra tions in d issol ved form of parent and daughter.
Thisequation can be rewritten as follows:
=

AD[D]
ApLpJ

4-2

where Ap and AD are the radioac tivi ties of parent and daughter.
As would be expected if the scavenging constant is zero, theradioactivities of parent and daughter will be identical.
On theother hand, if the scavenging probability is much greater than the
decay probabili ty, then the radioac tivi ty of the daughter isotope
will be much smaller than that of the parent.
Equation 4-2 can be rewritten as follows:

4-3f*Note that like the constants for radiodecay, the scavenging
constants have the units of inverse time.
It represents thefraction of the material scavenged from the water in a given unitof time.

171

or

4-4
where
11/2 andt1/2 are,respectively,
the half-lives forscavenging and for radioactive decay.
As can be seen from thisequation, if the activity of the daughter is one-half that of theparent, then the half-time for scavenging must be equal to thehalf-time for radiodecay.
Of course f (and 11/2) will vary fromelement to element.
The big picture of the resul ts obtained from measurements ofthe daughter-to-parent activity ratio for major oceanic watertypes is shown in table 4-4.
The main thing to be seen is thateverywhere in the sea there are processes at work which are capable of removing reac tive ions from the water column on a remarkably short time scale!
The apparent* scavenging times obtained
from equation 4-4 range from several decades for deep waters, toseveral months for open ocean surface waters, to several days forestuarine waters.
This progression is clearly of interest in thatit corresponds to the progression of particle concentrations in
the water.
If particles are responsible for the removal of thereactive sUbstances" then the removal probability would be expected to rise with particle content.
One might ask whether we can assume the missing daughter
radioactivity to be a measure of removal of the reactive ion fromthe water mass or whether this is merely the time for removal onto
particles which remain suspended in the water mass.
In our equations we designate the radioactivity of the daughter to be that insolu tion.
Ac tually the measurements of the isotopes of thoriwn
are generally made on the total samples (i. e., on d issol ved plus
particulate). Hence these results yield true removal times rather
than just particle uptake times.
For some of the lead and forsome of the thorium isotope measurements care has been taken to
separately es tima te the particula te frac tion.
These resul ts showthat in the case of open ocean samples more than 90% of the activity is in dissolved form (see figure 4-2 for an example). The implication from this finding is that the residence tLme for particles wi thin the wa ter mass must be considerably smaller than theresidence time for the reactive metals.
Hence the scavenging
times given here can be taken as removal times from the water
mass.

THORIUM ISOTOPES IN THE SEA
Having viewed the big picture, let us look at some of thedetails.
We will begin with the results on the three isotopes ofthorium (230 Th , 228 Th and 234 Th ).
Due to their widely differing
half-lives, each shows its greatest· sensitivity in a different*We say apparent because as we shall see, the assumption that theradiodecay of the parent is the sole source of the daughter isotope is not always valid.

172

Table 4-4. Typical activity ratios for daughter-parent pairs invarious water types.
Estuaries

CoastalSurface Ocean

Deep Sea

21 0pb/ 226 Ra->1*

0.4-1.0

23 0Th! 23 4 U--<3x 10- 5

3x10- 4

228Th/228 Ra

0.01

0.05

0.2

0.5-1.0

234 Th/ 23 8U

0.2

0.6

>0.9

- 1

23 ip a/ 23 5 U

-

-

-

2x 10- 3

210 po /2i0 pb

-

-

0.5

1.0

*Although 210 pb is being removed from surface water by particles,
it has an additional source. Radon atomsoescaping t02r~e atmosphere from continental soils decay to 21 Pb. ThesePb atoms
are incorporated into aerosols and are brought back to the earth's
surface by rain and aerosol impact. The flux of these atoms to
the sea surface exceeds b~ about a factor of 10 the in situ production by radiodecay of 26 Ra in the upper 200 meters of theocean.

173

SOLUBLE 210 pb • 226 Ra (dpm/IOOkg)

o

4

8

12

16

20

4

5

24

o

•
_

2

E
..:liC........

::c
I- 3

a.
w
o

4

5

o

I

2

3

PARTICULATE 210 Pb {dpm/lOOkg seawater)
Figure 4-2. Measurements of soluble 210 pb (circles) and particulate 210 pb (triangles) as a function of depth in the Antarctic
Ocean. The concentration of 226Ra is also shown (crosses).
The
concentra tion of 210 pb in particula te form averages only 0.5 dpm
per 100 kg of seawater, while that of dissolved 210Pb averages
about 14 dpm per 100 kg of seawater.
Hence only about 3% of the
210 Pb resides on particles and 97% is in dissolved form. These
measurements were made by Somayajulu and Craig (148) at the
Scripps Institution of Oceanography.

174place in the ocean.
234 Th , with its half-life of only 24 days, ismost useful in coastal and estuarine waters where the removal ofthorium is exceedingly rapid.
In the open ocean the scaven~ing
times are everywhere large compared to the half-life of 23 Th.
Hence its activity will everywhere be very close to that of its
parent.
228 Th , with a half-life of 1.9 years, is ideally suitedfor open ocean surface waters where the removal times are on theorder of one-half year, but too long for es tuaries where thescavenging time for thorium is on the order of days.
2
Shown in figure 4-3 is an example of the progression of
28 Th /228 Ra ratios away from a coast line (130).
Note the marked
decrease as the coast is approached.
The same data are plottedagainst salinity in figure 4-4 (due to the input of river water,
salini ty decreases from the open ocean to the coast).
As can beseen, the 228 Th /228 Ra ratio decreases with decreasing salinity.
The message is quite clear:
the scavenging time is much shorterin coastal water than in open ocean surface water.
It is of interest to compare results obtained from 228 Th withthose from 234 Th .
As can be seen from the graph in figure 4-5, in
those waters which have high enough scavenging rates to permit the
deficiency of 234 Th to be measured, the two isotopes yield reasonably concordant results (138).
Shown in figure 4-6 is a histogram of the 228 Th /228 Ra activity ratios for surface ocean water.
The corresponding scavenging
times are also shown.
The average half scavenging time is about
0.3 years and, as can be seen, the range is remarkably small. In
comparing these times with the longer ones we shall find for deep
waters, it must be remembered that surface waters are the habitat
for all plants and most animals.
In their quest for the criticalnutrient, phosphorus, plants may inadvertently take up reactive
metals.
These metals could then be passed along to the animals
and there be rejected in fecal pellets.
In this connection, it isinteres ting to note tha t the apparent sca venging time of PO 4 inthe surface ocean 1s of the same order as that of thorium.*
One of the most spectacular sets of findings in this area ofthe research comes from the measurements by Nozaki and his coworkers in Japan of the dissolved 230 Th concentration in verylarge sea water samples (135) and the measurements by Krishnaswami
and his colleagues in India of the 23 °Th bound onto particles
filtered from the deep sea (137).
Nozaki's results for a profilein the northwestern Pacific are shown in figure 4-7. The activity
increases at a nearly constant rate down the water colums.
The
corresponding apparent half-removal times (calculated from equation 4-4) as shown along the bottom of the diagram range from lessthan a few years in surface water to about 50 years for bottom* As sh?wn in chapter .1,

the removal time of PO 4 from the upper
ocean (l.e., surface m1xed layer plus one half of thermocline) is
about five years. However, as the rates of photosynthesis exceed
the rates of rain of organic matter from the mid-thermocline tothe deep sea by roughly a factor of 10, the scavenging time forphosphorus by plants must be about one half year.
The difference
betw:en.the rate of photosynthesis and the rate of rain of organic
debr1S 1S the result of nutrient recycling within the upper ocean.

175

LONGITUDE oW
41 ~__...,..~--::~a;:7c3===-~~==,7""'1'2 ~-=_-r-_.,

z
°w 40
Q
~

....

....
<l:
....J

391----+----:-r---++~~~---t-z°

lJJ 40
Q
~

....
....
<t

...J

.. 091Figure 4-3- The upper map shows the distribution of 228 Ra concentrations (circled values) and of 228 Th concentrations in units ofdisintegrations per minute ~~r 100 kg of water and the lower map
the distribution of 228 Th /2 Ra activity ratios in coastal waters
adjacent to New York City.
These measurements were made by Feely
and his colleagues at the Lamont-Doherty Geological Observatory ofColumbia University (130).

176

20
0'

.¥

0
0
......
E~•-. •••••.#•• •• • •• • •
• "
•

15

..

10

~
0

0::

CD

N
N

-

5
0

Cl

31

0
0

0.6

~

0.4

::::
e
~

•

0.2

•
-0.15
0-;
-

N
N

0

35

•
•
•

•

• • ··~,r•.,.\-.,. .

.s::.

,r-

34

33

32

~

••

•

0

32

0
'-

34

33

>-

0

0

0.10

0

0::
to

N
N

0.05

......
.s::.

I-

CD

~

0

•31

..
"
.
,:e-••
\

f

32

~

33

•
•
34

35

SALINITY (°/0)

Figure 4-4. Concentra tions of 22 BRa (upper) and of 22 8Th (middle)
plotted against salinity in coastal samples taken adjacent to New
York Ci ty (see figure 4-3 for map).
The ac ti vi ty ra tio of these
two isotopes is also plotted against salinity (lower). Due to the
input of river water, the waters closest to the coast have thelowes t salini ty.
Hence the increase in salini t.v is a measure of'the effective distance from the shore.
As 2'28 Ra is added tosurface waters from shelf sediment, the concentration of thisisotope decreases away from the shore.
As plant biomass andcontinental detri tus burdens of the wa ter decrease away from thecoast, the scavengi~ probability for 228 Th decreases.
Hence theratio of 228 Th to 2 Ra increases with increasing salinity (hencewith increasing effective distance from the shore).

177

~

0.7

0

....0
~

0.6

....
....u> 0.5~•0.........

:J

0.4

~

".c.

.-

V

~

0.3

•

•

0.2

C\J

0.1

•

•
•
•

, •

<X)

C\J

•
••

•

0

0.01

0.02 0.03 0.04 0.05 0.06 0.07 0.08 0.09
228 Th /228Ra (octi vity ra tio)

Figure 4-5. The activit~ ratio, 234 Th /238 U, plotted against theactivity ratio, 228 Th /22 Ra, for coastal waters.
The curve forconcordance between the scavenging times calculated from these two
ratios is also shown.
These measurements were made by Kaufman at
Yale University (138).

178

HALF SCAVENGING TIME (yrs)

.2 I
.48
.8 I
I. 27
I. 9 r25 .........- .....--............- ..........- ......- .....--.-..
.00

P = PACI FIC
I = INDIAN
A = ATLANTIC

en 20
W

~

C1~

«
en

15

lJ..

0

a::
w 10

CD
~
~

z

5

.1

.2

.3

.4

.5

228Th ACTIVITy/228 Ra ACTIVITY

Figure 4-6. Histogram showing the 228 Th/228 Ra ratios for open
ocean surface waters from the Atlantic, Pacific, and Indian
Oceans.
The corresponding half scavenging times as calculated
from equation 4-4 are also given.
These measurements were made at
the Lamont-Doherty Geological Observatory (125).

179

0.0

O. •

,

\.

I

.-::r::a.. 3
w

o 4

0

°.

.

I

•

I
...-.

• \

51-

6'

(PARTICULATE)

.00

,

.01

.

,

.02

N W PAC I F I CEoX........

.\
.\
.\

~

...-

.00

,

'\

-E 2

230Th ACTIVITY (dpm/IOO kg)

230 Th ACTIVITY (dpm/IOO kg)(DISSOLVED + PARTICULATE)
O. I
0.2
0.3

:I:

I-

2

•

I

0

3

a..

.~

w
0

0

4

I/IOth DISSOLVED
ACTIVITY

1 I•

"

~

5

~

,

,

25

50

INDIAN OCEAN

o PA~IFIC OCEAN,

75°APPARENT HALF SCAVENGING TIME

6

(yrs)
F1.gure 4-7. In the left-hand panel is a plot of the acti vi ty of the total 23 °Th insea water as a function of depth in the northwestern Pacific Ocean.
For reference
the activity of 234 U (the parent of 230 Th ) is 280 dpm/100 kg. These measurements
were made by Nozaki in Japan (135).
In the right-hand panel are measurements madeby Krishnaswami and his colleagues from the Applied Physics Laboratory in Amadabad,
India, on particulate matter filtered by an in situ apparatus from waters in the
Indian and North Pacific Oceans (137).
--

waters.
Krishnaswami' s results for 230 Th on particulate matterare also shown in this figure.
As can be seen, they have a depth
pattern similar to that for dissolved 230 Th , but per unit volume
of water carry only one-tenth as much 230 Th . These results reveal
an important aspect of metal geochemistry.
The adsorption ofthorium onto particulate matter must be reversible, and for the
deep waters of the open Pacific Ocean, the equilibrium distribution coefficient between water amd particles must be about 10.
To see why this is the case, we mus t fi rs t unders tand tha t,since more than 99.9% of the 23 °Th produced in the ocean reaches
the sediment, the particles falling through a given horizon in thesea must carry an amount of 230 Th proportional to that produced in
the overlying water column.
Thus the amount of 230Th carried byparticles must increase linearly with depth.
In this regard
Krishnaswami's findings confirm expectation.
If reversible equilibrium is maintained between the Z30 Th concentration on the falling particles and that in the water and if the concentration of~articles remains
constant wi th depth, then the concentration of
30 Th in the water must increase linearly with depth as does thaton particles.
The concentration of 230Th on the particles, Cp , will be
determined by the requirement that the flux of 230 Th atoms throughany horizon equals the production rate of 230 Th atoms in theoverlying water column.
Hence:
c.pS

Pz

4-5

Cp =

Pz atoms
g
S

4-6

or

where S is the flux of particulate matter (g/cm 2yr ), P is the
production rate of 230Th (atoms/cm 3yr), and z is the depth (cm).
The concentration of 230 Th in the water (C w ) is given by:
C w = KC P

4-7

where K is the distribution coefficient (i.e., the ratio ofthorium in a gram of wa ter to that on a gram of particles).
Combining equations 4-6 and 4-7, we obtain:
KPz
C = - a t oms
4-8
w

S

g

At any glven place in the sea the concentrations of 230 Th in thewater will build up to that point where the particles passing
through the water carry an amount of 230 Th equal to the productionin the overlying water column.
Once this concentration has beenachieved, the system will be at steady state.
230Th will be removed as fast as it is generated.
If the average settling veloci ty of particles is taken to be everywhere the same, then theconcentration of particles in the water will be proportion~oto
the particle flux.
In this situation the concentration ofTh
in the water will be inversely proportional to the particle
concentration.

181

Nazaki's resul ts also suggest that recycling of 230 Th fromthe sea floor is unimportant. Were it important, one would expect
the slope of the 20Th activity versus depth relationship to
steepen near the bottom. No such change is seen.
Based on the apparent scavenging time of several tens ofyears calculated from Nozaki' s 23 °Th data, it might be expectedthat the 228 Th/228 Ra activity ratio in deep water would be close
to unity. Measurements of these isotopes on samples of deep watercollected as part of the GEOSECS program are shown in figure 4-8.
At all these stations the 228 Ra content decreases away from thebottom.
As 228 Ra (half-life 5.7 years) is supplied to the sea
entirely from sediments, its concentration should be highest near
the sea floor.
Of in teres t to us here is the difference between
the acti vi ty of 22 8Ra (the parent) and 22 8Th (the daughter).
As
can be seen from the figure 4-8, the evidence is mixed.
At someof the GEOSECS stations (28, 34, 58 and 263) the activities ofthese two isotopes a~ree within the experimental error. At others(32,61 and 2142 the 28 Th activity seems to be significantly lower
than that of 228 Ra .
It can be shown that if the contribution of 228Th produced
from the 22 8Ra in the upper ocean is neglected, and if the distribution of 228Ra above the sea floor follows the relationship:

4-9
where z is the distance above the sea floor and h is the characteristic height above the bottom attained bl, 228 Ra atoms before
undergoing radiodecay, then the ratio of 22 Th to 228Ra will begiven by the expression:
1SK(1 +Kq) Ah

+ 1

4-10

S is the flux of particulate matter (g/cm 2yr), K is the distribution coefficient of thorium between particulate matter and water(atoms per g particulate matter/atoms per g H20), q is the concentration of particulate matter in the water, and A the de cay constant of 22ffTh .
The combined work of Nozaki et ale (135) and Krishnaswami etale (137) shows that the product Kq is about 0.1. Hence equation
4=10 can b~ simplified as follows:1

4-11

SK + 1
Ah

Taking S to be 1 x 10- 3 g/cm 2yr, K to be 10- 7 ,to be 7 x 104 cm (i.e., 700 m):

182

A

to be 0.363 and h2...-..-.....

~GEOSECS
STN
28
ATLANTIC

3
+

+

I

'-4

a..
W
o

+ -0----<r-

+

--0~

5

+
++

6---

o 0.4 0.8

3~~-0----0--0.4 0.8 1.2

1.6-'0.4 0.8~

r.-I

0.4 0.8- -. . . .

0.4 0.8

228 Ra and 228 Th (dpm/IOO kg)
...,.
-.-~
-.~

......

GEOSECS
STN
203
PACIFIC

_

=.E4

I
I-

a..

0

-I-

0

+

0

~5

0.4 0.8

0.4 0.8

0.4 0.8

0.4 0.8

0.4 0.8

0.4 0.8

228 Ra and 228 Th (dpm/IOOkg)
4-8. Plots of 228 Th activity (crosses) and of 228 Ra activity (circles) as a function of depth in the deep waters of thewestern Atlantic Ocean (upper) and in the Pacific Ocean (lower)
The bottom depth is shown at each sta tion.
These measurements
were made as part of the GEOSECS program by Herbert Feely and hiscolleagues at the Lamont-Doherty Geological Observatory.
Figure

183

1

1x10-

3

1x10 7

X

0.363 x 7 X 10 4

= 0.72+ 1

The exact value will of course depend on the local particle flux(the larger the particle flux, the smaller the ratio) and on thelocal characteristic height attained by 228 Ra atom (the greater
the height, the closer the ratio will be to unity).
The point tobe made is that deficiencies of 228 Th activit~ in bottom water are
not inconsistent with the observations on 23 Th.
Rather, in situations where daughters are carried into a water mass on particles
raining down the water column, use of the scavenging times obtained using equation 4-4 can be very misleading.
Before leaving thorium,
one other set of datf28mus~28bementioned.
Shown in figure 4-9 are profiles of theThlRaactivity ratio in the upper thermocline of the Atlantic Ocean(132) . These ratios show an increase wi th depth reaching values
greater than uni ty at a few hundred meters.
These resul ts canonly be explained if the 228 Th atoms removed from the surface
water are released from their carrier particles in the upper
thermocline.
The most likely scenario for such release would bethat the 228 Th atoms born in surface water are inadvertently
incorporated by plants during their quest for phosphorus.
Theanimals which eat these plants incorporate these 228 Th atoms into
fecal pellets.
The ingredients of these pellets are eventually
released to the water when the orEPnanic material they contain iseaten.
Once released, these 22 Th atoms would of course bevulnerable to uptake onto the particles present in the upper
thermocline.
So yet another complication enters our story.
Uptake by living organisms coupled with release during respiration
appears to compete with inorganic absorption at least in the upper
water column.

PROTACTINIUM-231 TO THORIUM-230 ACTIVITY RATIOS
Although measurements of dissolved 231 pa activity versus
depth in the wa te r column have yet to be made, we do have some
information about the removal of 231 pa from the sea which adds toour story.
This information comes from sediments and s1dimenttrap material.
The activity ratios of 231 pa to that of
30 Th insediments from well away from continental margins are lower thanthose expeeted from the ratio of the production rates of these two
isotopes in sea water. * As the integrated excess 230 Th activity*T~•
f 235U238.
,
21 e ratio 0
tf35.._ U atoms 19 everyw~ere constant I~ ~he solar system.
All materials in~~\\ding sea water have 138
atoms for each-U a~~JB'
Due to recoil damage, the3", atoms generate~ ~y the decay of
11 in mineral can become
separated fr~m the parent.,?' In the sea, for e~a~ple. the ac~ivity ratio 3 U activity to 23'\r activity is l.15 (27).
Because uranium atoms reslde.ln the .sea, for many mlx~~ cycles (}~~ •• for2~,ny times ~~~O years). this ratio is everywhere
the same. Hence the production ratio (In atoms) ofTh (fromU) toPa (fromU) is given by:R-

prod -

1.15 ).238U N

238>.235 U N23S U

U

..

lO

1.15x1.55xlO9.85xlO- 10

138

= 25

1

The decay rate rat io for a batch of newly produced 23llrh and 23"8 is given by:_ >.230 Th

Rdecay - A~ Rprod

9.22xlO- 6= 2.13xIO- S 25.0 = 10.8

184

a
100

E

....... 200
I

l-

n.

w 300
0

400

0

20

40

60S

LATITUDE
0o r--)C-)C00(7)a;N~V(7) 0'1 0'1

t-A

oo

(J'1

~

i

a8
O.2&~

loot : ~ :-:-~-~
(JI

x

~

x

=.: 'f&~
• x

200

N
0

In.

x

t-

.

~-::-J?
~•~!!2~t:---STATIONNUMBER

.,

~

It)

~ 300
400
60S

40

20

0

20

40N

LATITUDE
Figure 4-9. Sections in the western (upper) and eastern (lower) Atlantic of theactivity ratio of 228 Th to 228 Ra in the upper 400 meters of the water column based on
results from the GEOSECS program.
As can be seen at every station, this activity
ratio increases with water depth, reaching values greater than unity at depths
ranging from about 100 meters in the equatorial zone to 200 to 300 meters in thetemperate zones.
These resul ts were obtained by Feely and his colleagues at theLamont-Doherty Geological Observatory (132).

in these sediments matches reasonably well the decay rate of 234 U
in the overlying water column, this anomaly a~pears to be theresult of a lower than expected rain rate of 23 Pa.
If so, thenthe 231 pa produced in the water column above these sites must movela terally in the ocean and undergo removal elsewhere on the sea
floor.
This "elsewhere" has been found recently by De Master at
Yale (129) and by Mangini at Heidelberg (personal communication).
Their measurements show that sediments forming beneath areas of
upwelling in the Antarctic and along continental margins have muchhigher than expected 231 pa to 230 Th ratios. While enough data has
not been obtained to show that the excess 231 pa rain in theseareas compensates for the deficient rain of this isotope in the
open ocean, it seems likely that this is the case.
More direct evidence for this phenomenon is available fromsediment trap data.
Anderson and Bacon at Woods Hole haveanalyzed material from Honjo' straps (134) and have shown ~hattraps deploy-ed in the open ocean show a deficiency of 2 lpa
relative to 230 Th (see figure 4-10 for results and figure 4-11 forlocations at which the traps were deployed).
On the other hand,t raps deployed close to Panama gi ve jus t the 0roosi te re~~l t.
They accumulated an anomalously high ratio of
31 pa to 2 Th.
Care must be taken when converting sediment trap results toabsolute rain rates.
The efficiency of these traps varies with
their dimensions and shape, wi th particle size, and wi th current
velocity.
No one has as yet succeeded in establishing a means tocalibrate these devices.
Nevertheless, if the efficiency of thetraps is assumed to have been the same in all four locations, thenthe absolute rain rate of both 23 °Th and 231 pa must be considerable higher near Panama than for the open ocean sites.
These observations can be explained by the combination of
rapid lateral mixing wi thin the sea and particle concentration
dependent scavenging rates.
This phenomenon can be best understood by considering the extreme case where lateral mixin;p isinfinitely rapid.
Under such ci rcums tances the number of
30 Th
and 231 pa atoms reaching a gi ven place on the sea floor would be
proportional to the number of particles reaching the bottom.
Thus, if an area of the sea floor near a continental margin
recei ved 10 times more particles than an area of the sea floor
well away from the margin, it would also recei ve 10 times more
230 Th and 231 pa • This would lead to a higher than expected sediment inventory of these isotopes near the continental margin and alower than, expected inventory in the area away from the margin.
If our hypothetical sea had two types of area, 10% of the highflux type and 90% of the low-flux type, just about half of theseisotopes would end up beneath each type of regime.
The sediment
inventory would be about half the expected value in the 90% of thesea f I 00 r re c e i vi ng th e low flu x , and abo u t 5 time s the exp e c ted
value in 10% of the sea floor area receiving the high flux.
The distributions of 230 Th and 231pa on the sea floor and in
sediment traps can be explained by postulatin~ a shorter wa ter
column scavenging time for 230 Th than for 23 Pa, and by postulating that the average of the 230 Th and 231 pa scavenging times iscomparable to the time required for lateral homogenization along
isopycnal surfaces in a major ocean basin.
In such a si tuation

186

UNSUPPORTED 230Th (dpm/em 2103yr)
00
2
46
e 10 12 14 16

UNSUPPORTED 231 Po (dpm/em 2 r0 3 yr)
6'.0 0.2 0.4 0.6 0.8 1.0 1.2 1.4• SITE S• SITE S

r::J SITE E
x SITE P

l:J SITE E
x SITE P

2

~ 3
a.
ILl
o

UNSUPPORTED 230Th/231po (oell,,"y rolla)
00
10
20
30
40
• SITE S['] SITE E
I 1-)( SITE P

2

~

l:J

)(

.

2

E

x

~

~ 3

x

z 3-

.

~

a.

Q.

l&.I
0

•

[']

4

•

W
0

r::J

•

l:J

4x

x

~EXP. FROM-

234U/23~u

~

~~

l:J

IN SEA WATER

.

x

6

UNSUPPORTED 230Th (dpm/cm 2 10 3yrl
~
10
I~
20
2

UNSUPPORTED

2

231

2
Po (dlm/cm 10yr)

6

)(

•

I

UNSUPPORTED 230Th/231F'a (octi"lty rolio)
2
4
6
10
12

3

oc

• SITE I

SITE I

-

l:J

e

•

SITE I

•
II-

e

E

:5
%

~

E

EXP. FROM WATERCOLUMN PRODUCTION

~

0

:3

4

•

~

• •

•

3

•

4

••

0.
ILl
0

•

•

4

-

IN SEAWATER

z 2-

•

•

w

-

234U123~u

~

2

Q.

• EXP. FROM~

.

~

•

I

I

-

.

Figure 4-10. 230 Th and 231pa fluxes as determined from sedimentt rap deployments at various loca tions and depths in the ocean.
The upper panel shows the results from three open ocean sites (forlocations, see figure 4-11).
The lower panel shows the resultsfrom a sit"e close to the continental margin.
The solid lines
represent the results expected if the traps were collecting theseisotopes at the rates they are produced in the overlying watercolumn. These results were obtained by the geochemistry group at
Woods Hole Oceanographic on samples from Honjo I s sediment traps(134). By unsupported is meant the activity excess over that for
the respective ~arent isotopes (234 U in the case of 230 Th and 235 U
in the case of 31 pa ) in the trap.

187

Figure 11-11. Map showing the sites of sediment traps de~loyed by
Honj 0 and his Woods Hole colleagues.
The 23 °Th and 2 ipa da ta
concentrations in the material caught in these traps are summarized in figure 4-10.
r--1IIfI""'l~""-~-------",,,-"'-""'''''--''''''-''''''--'''-''!''''''I'''''''''''60°...-.. .0#'.:..

J• • •r .

Table 4-5. Ratio of 230 T h to 231 pa removal time as determined fromtheir activity ratios on Mn02-coated nitex. Results obtained byAnderson and his colleagues at Woods Hole Oceanographic (134).
Depth
km

A230 Th
A23I pa

T231 pa
T230 Th

Site*

5.8
4.2
4.2
4.2
4.2
4.2

3. 4
5.3
5.2
6.1
5.7
3.9

3.2
2.1
2.1
1.8
1.9
2.8

P

*see figure 4-11.

188

I
I
I
I
I

both
isotopeswouldshow
higherrain
rates(and sediment
inventories) in areas of high particle flux than in areas of low
particle flux.
The contrast, however, would be greater in thecase of 23 1Pa than in the case of 23 °Th.
This would gi ve rise to
the higher than theoretical 23 °Th/ 2 3 1Pa ra tios seen in the deepsea and the lower than theoretical ratios found in areas of highparticulate rain rate (i.e., along continental margins and in the
Antarctic).
Evidence exists
for
this
difference in removal times.
Anderson of Woods Hole deployed manganese oxide coated passive
absorbers on sediment trap moorings (134).
Assuming that theseabsorbers have the same uptake efficiencies for dissol ved 23 °Th
and 231 pa , the ratio of the activities of these two isotopes on
the absorber should be equal to the ratio of their activities in
sea water.
Were the scavenging times of these isoto~rrs by marine
particulates identical, then the activity ratio of 2 Th to 231Pa
in dissolved form should be equal to 10.8 (i.e., the decay rate
ratio of newly formed 230 Th and 321 Pa ). As can be seen from table
4-5, the ratios observed by Anderson on his passive ion absor~~rs
ranged from 3.4 to 6.1.
The suggestion is then that the 2 Pa
resides 1.8 to 3.2 times longer in sea water than does 230 Th .
This result should of course be confirmed by direct measurementson large-volume water samples.
A similar result is obtained if manganese nodules are assumed
to absorb 231 Pa and 230 Th in proportion to their concentrations in
sea water.
During the mid-1960 r s Sackett showed that the 23 °Th/
231Pa ratio in nodules averages about half that expected from theratio of the production rates of these isotopes in the sea (220).
Again the implication is that 231 Pa atoms reside in the water
about twice as long as do 230 Th atoms.
Data on 230 Th to 231 Pa activity ratios in sediment traps andin core tops are summarized in figure 4-12.
As can be seen, theratio of these isotopes is inversely related to the particle content of the overl ing water column. Where the particle content islow, the 230 Th/23 1Pa ratio is high. Where the particle content ishigh, the ratio is low.

THE DISTRIBUTION OF LEAD-210
Let us next consider the isotope 21 °Pb.
Its distribution in
the sea is complicated by the fact that it is supplied via atmospheric "fallout" as well as by the radioactive decay of the 226Ra
in sea water (139).
The atmospheric supply is derived from radon
atoms which escape to the atmosphere from soils and then decay toform 21 0Pb.
These atoms become at tached to aerosols which carrythem back to the land or sea surface.
While the production rateof 210Pb atoms wi thin the water column can be exactly determined,
the input rate of 210 Pb atoms from the atmosphere to the sea surface is only roughly known.
A summary of 210 pb fallout rates
measured over land and of water column production rates in thevarious oceans is given in table 4-6.
Assuming that the fallout
rate of 210 Pb over the sea averages roughly the same as it does
over land, then the atmosphere contributes about 20% of the amountof 210 Pb in the sea.
While not greatly altering the situation in

189

........

•

.0-

~

c

~

~

~

....>

•

()

c
.....,-~PRODUCT OF
PART. CONC. 230 Th /23I pa

r t)

C\J

'"
....

•

~

0

EQUAL TO A CONSTANT

rt)

C\J

EXP. FROM

0234 U /235 U

W

....
0::

IN SEA WATER

0

C-

o.
::J

en

•

z

~

00

10
20
30
40
50
PARTICLE CONCENTRATION (jLg/l)

Figure 4-12. Unsupported 230 Th to 23I pa activi ty ratios in sediment core tops, in sedlinent trap material and in particles filtered from the water as a function of the average particulate concentra tion in the overly ing wa ter column for a wide range of ocean
sites.
These data were compiled by Bob Anderson (134).

190

Table 4-6. Measurements
production rates.

of

210Pb fallout

rates and water column

Fallout Rates
Location(Ref. )

Material

Fallout Rate
atoms/cm 2min

Greenland
Antarctica
Mt. Olympus, WA.
C.har Lake,
Canadian Arctic
Reindeer Lake
Sask .. , Canada
E.L.A., Kenora, Can.
Lake Superior
Lake Michigan
Lake Ontario
Lake Shinhi, Japan
Lake Leman, Switz.
Lake Pavin, France
Fayetteville, Ark.
fv1aryland
C.onnecticut
Pennsylvania
New Mexico
C.olorado
South Africa(140a)(140a)(140a)(147a)Ice
Ice
Ice
Lake Sediments

3.6
0.2
5.4
3.0

(147a)

Lake Sediments

9(147a)(142a)(146a)(147a)(147b)(142a)(142a)(145a)(140 )(155)(155 )(155)(153 )(154)

Lake Sediments
Lake Sediments
Lake Sediments
Lake Sediments
Lake Sediments
Lake Sediments
Lake Sediments
Rain
Soil
Soil
Soil
Soil
Soil
Soil

17
13
21
19
25
14
41
32
39
26
33
65
52
21

Production Rates
Location

Production t
atoms/cm 2rnin

North Atlantic(Vies tern Bas in)North Atlantic(Eas tern Basin)
South Atlantic(Wes tern Basin)
South Atlantic( Ea s tern Bas in )Antarctic
Central So. PacificCentral No. Pacific

Fallout*
atoms/crn 2minTotal
a toms/ cm 2min

54

24

78

60

24

84

66

12

78

72

12

84

78
90
120

< 5
12
24

~

80
102
144

tFor a 4000-rneter water column.
*The northern hemisphere values are based on the data in thefirst half of this table; the southern hemisphere values are
based on the assumption that the fallout rate in the southernhemisphere is one-half that in the northern hemisphere.

191the deep sea, the airborne component complicates the situation in
the upper water column.
A typical depth profile for 210 pb activity is shown in figure
4-13.
In the upper water column the acti vi ty of 210 pb exceedsthat of 226 Ra (i.e., atmospheric fallout more than compensates forthe 210 pb removed onto particulate matter). At mid-depths th~l~wO
isotopes have similar activities.
In bottom waters thePbactivity is only about half the 226 Ra activity (clear evidence forthe removal of L10Pb onto particles).
Let us start our analysis of the 210 Pb distribution with waters below 1500 meters where the contribution of atmospheric fallout is relati vely small.
As shown in figure 4-14, profiles of
210 pb from various locations in the North Pacific and of the Antarctic all show the same basic trend.
The activity ratio dec reases toward the bottom.
As shown in figure 4-15, this si tuation does not change dramatically in the near-bottom layer (157).
As do all the waters in the lower part of the water column in the
North Pacific, these near-bottom samples hav~ a
210 Pb /226 Ra
activity ratio close to one-half. Thus, unlike 20 Th , which shows
an apparent half-s cavenging time increasinA from 20 to 50 years
over the depth range 2500 to 5000 meters,°Pb shows an apparent
half-scavenging time decreasing from more than 100 years at
mid-depths to 15 years near the bottom.
In this regard, it would be interesting to know what scavenging time in upper waters is.
Here we run into the difficulty
associa ted wi th the a tmospheric input ~As we do not know wi th
sufficient accuracy the delivery rate of these atoms, it is notpossible to give a firm estimate.
We do have a few clues however.
The map in figure 4-16 shows the distribution of 21 t Pb in
surface waters in the Pacific Ocean (151). The difference between
the hemispheres reflects the higher land area in the northernhemisphere, which leads to a greater total radon release and, inturn, toa greater mean 210 Pb fallout rate.
The important feature of this diagram is rather the drop in acti vi ty toward theocean margins.
This decrease very likely reflects the higher
particle concentrations in coastal waters.
Thus we have at least
quali tative evidence that 210 Pb is being removed in significant
amounts from surface water on a time scale of its half-life (i.e.,
22 years).
Another clue comes through a comparison of the upper watercolumn inventory of 210 Pb activity with the rate of fallout asmeasured over the continent.
In fi9.lbre 4-17 are shown the composite of 'a series of profiles of 2 Pb and 226 Ra for the upper
600 meters of the North Pacific Ocean.
Note that the excess
extends to about 600 meters.
As this is the depth to which bombproduced tritium had penetrated in the decade between its addition(~1964)and the GEOSECS survey (1974), physical mixing of theatmospherically deri ved 210 pb could be entirely responsible forits penetration (i.e., uptake onto, followed by [e~ease from,particles need not be called upon). The integrated 1 Pb activity
in this layer is 11 dpm/cm 2 • Based on continental measurements,
the flux of 210 Pb atoms onto the sea surface should lie in therange 30 ±1 0 atoms/ em 2 minu te.
Adding to this the 6 atoms/ cm 2
minute produced by the 226 Ra dissolved in the upper 600 meters, we

192

ACTIVITY (dpm/IOOkg)
0

~

E

20

10

~0 c9

0

0

jJ

•
0

• •
•
•

2

30

0
0
0
0

• • <=-.::t:.

0

~

.-

:I: 3
0W
0 4

GEOSECS STN 214
32°N 177°W

226
210

Ra 0
pb

•

5

6

•

•
•

•

•

• •

0
0
0
0

0

Figure 4-13- 22 6 Ra and 210 pb ac tivi ties versus depth at GEOSECS
station 214 in the central North Pacific.
The 226 Ra data wereobtained by Chup§ of the Scripps Institution for Oceanography(118).
The 21 Pb data were obtained by Turekian and hiscolleagues at Yale University (158).

193

ACTIVITY (dpm/lOOkgl

ACTIVITY (dpmIlOOkg)
2030_~----.~"T"T

.

40

0r-

.~
.

:~o

\

n.
w
a

•

3 f.-

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O.B

1.0

90yr

>400yr

Q.8

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-

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•

6

3

3

n.4

n.4

3 LAT. LONG.210Pb226Ra

,

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a.. 4
w
a

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0.4

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90yr

30-.. 160·W.

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lTTT"

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5 LAT. lDNG. 2IOPb 226Ra

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@

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30

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ACTIVITY (dpmlIOOkg)
10
20

2 f.-

2

@

210 pb •,~-WEDDELL SEA

6'--~

,now.

1.0

E)

•

B'N

0

Ii)

226 Ro

E

-

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a!

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2

·@l
•

•

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a

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2

ACTIVITY (dpm/IOOkg)
10
20
30

o

IT"l0.-----.~2T"0~-~3T"0-.,(;)-

LAT. LONG. 210pb 226 Ra
60'S
6'---I---~--Io-_.L&-_ _..I-II

6

O'

,•

0

•

•

Figure 4-14. 226 Ra (open symbols) and 210 Pb (closed symbols) activities for water samples taken at various depths in the deep
Pacific and Antarctic Oceans.
In all cases the deficiency of
210 Pb is greater in bottom water than in mid-depth water.
Thesolid curves represent various activity ratios for 210 pb and
The 1.0 curve is a best fit to the 226 Ra data.
226 Ra .
The scavenging times (as calculated from e~uation 4-4) corresponding tothese ratios are also given.
The 22 Ra results are those of Chungat Scripps Ins ti tu tion of Oceanography (117, 118,160) .
The 21 °pb
results are from the Yale group (158), Hakodate Japan group (152),
and the Scripps Institution for Oceanography group (160).

194

210 pb

ACTIVITY (dpml/OOkg)

10

0.4

I

20

I

I

I

. I

30
I

•

28.5°N
145°E........-.

E
;0.3
e
~

-

-

I-I--

e

•

ID

I

w
~ 0.2 -

al

«

w

0.4

0.6

0. 8

1.0 -

15yr

33yr

90yr

>400yr

•

u

z
«

~ 0.1

-

o

o

I

•
.-

I

-

I

I

I

Figure 4-15. Near-bottom 210 pb results for a station in the northwest Pacific.
The 226 Ra contents were not measured but can be assumed to be the same as the values shown in figure 4-14.
thelines represent various 210 pb /226 Ra activity ratios.
The 21 Pb
results were obtained by the group in Hakodate, Japan (152).

195

210 pb

dpm/IOOkg

Figure 4-16. Map showing the dis tribtuion of 21 °pb concentrations(in dpm/IOO kg) in the surface waters of the Pacific Ocean. The
results were summarized by Nozaki, Thompson and Turekian (151).

196

ACT I V I T Y (dpmllOO kg).• ...•..
0• i.r;•

10

0
0.0

••

•
••
•
•• ••
•
III

0

0.2

0

0

,e "
0

E

•

:r:

0.4

I-

aD

w

o

a.
0

:'.

;;,~:e ••0 0

32°N 177°W
AS OF 1974
0.2

0.4

0@.
0

•

0.6

0.6

0.8

.

0

0.0

••
••••
• • ••

0
0·0

x

.~

30

0
•

226 Ra210 pb

4-17. The left hand panel shows the vertical dis tribu tionof 226 Ra and 210 pb activity in the north temperate Pacific Ocean.
The solid squares represent 210 Pb measurement made by the Japanese and the solid circles 210 pb measurements made as part of the
GEOSECS program.
Both sets show scatter beyond that seen in otheroceanographic data sets.
Whether this sea tter is the resul t ofproblems with the measurements or represents large inhomogeneities
in the 210 Pb distribution is not known. The open circles represent
226 Ra measurements made as part of the GEOSECS program.
The
right hand' panel shows the distribution of tritium with depth in
the center of this area as of late 1973. The tr-itiwn r-esults wereobtained by Ostlund as part of the GEOSECS program (54).
The
226 Ra results wer-e obtained by Chung of Scripps Institution of
Oceanography (118).
The 210 Pb resul ts were obtained by the Yalegroup, solid circles, (158) and by the Hakodate Japan group, solid
squares, (152).

Figure197

get a total input of 36 ± 10 atoms/cm 2 minute.
Thus, 25 ± 11
a toms/ cm 2 minute must be lost to raining particles.
If so, theaverage scavenging time for 210 Pb in the upper 600 meters of thewater column is about 10 years.*
There is one last war to get an estimate of the average watercolumn removal rate of 2 °Pb.
That is to use data froP3osedimenttraps.
Shown in table 4-7 are the ratios for 210 Pb to 2 Th atoms
caught in deep traps at sites E and S (155a).
In both cases about
two 210 pb atoms are caught for every 100 230 Th atoms.
Based onthe water column 234 U and 226 Ra distributions at these sites, it
can be shown that about five 210 Pb atoms are produced for each 100
230 Th atoms.
Adding in the 210 Pb atoms from atmospheric fallout,
the ratio rises to 6 210 Pb atoms per 100 23 °Th atoms.
We knowthat over 99% of the 230 Th atoms produced in the water column rain
to the bottom.
Hence these results suggest that 33% of the 210 pb
atoms rain to the bottom.
If so, the apparent scavenging time forlead is 45 years.
This resul t is consis tent wi th the mean watercolumn s ca venging time obtained from the 21 °pb to 226 Ra ratio in
wa te r samples.
In the same table are given the atom collection ratios for
234 Th and 230Th.
Only about six 234Th atoms are caught per 10,000
230 Th atoms.
As the two isoto£es are produced at nearly the samerate, about 99.94% of the 23 Th atoms must decay in the watercolumn!
This isotope has a half-life of only 0.066 years, hence
the corresponding half-scavenging time is about 100 years.
As
some of the 234 Th atoms undergo decay while riding toward thetraps on particles, the removal time from a gi ven level in theocean would be less than 100 years.
Again the result is consistent with that derived frora the 230 Th and 228 Th measurements.

THE DISTRIBUTION OF POLONIUM-210
The last of the parent-daughter pairs to be discussed here is
210 po _210 pb •As shown in figure 4-18, for waters beneath about
100 meters, the acti vi ties of the two isotopes are nearly equal.
As the half-life of 210 po is 0.38 years, this is to be expected if
the scavenging time for polonium in deeper waters is similar tothat for lead (i.e., several decades).
In surface waters from both the Atlantic and Pacific (see* Most workers in this field interpret the 210 pb data to yield ascavenging time for lead atoms in surface wa5er of about one
year~ This is the answer obtained if all the 21 Pb atoms reaching
the sea surface through atmospheric fallout are assumed to be
removed from surface water by scavenging.
This assumption neednot be valid.
As shown in figure 4-17, tritium (which must movewith the water) is carried to a depth of many hundreds of mete son a time scale of 10 years.
The mixing processes which carry 3H
atoms to these depths must, of course, also carry 210 pb atoms. A
more conservative interpretation of the data is to treat theentire upper water column and conclude that the average scavenging
time for lead over this entire interval is 7 years.
Only if the
scavenging could be shown to occur mainly in the surface mixed
layer would the one year scavenging time be appropriate.

198

Table 4-7. Measurements of uranium series nuclides in sediment
trap rna terial obtained by the \'Joods Hole Geochemistry Group on
samples from Honjo's traps (128).

Atom Ratio
in Traps
Prod. Ratio
in Water
Column**Relative
RemovalEfficiency
Apparent HalfScavenging
Time* (Years)

Site E

Site S

5086 meters

3694 meters5.5x10- 4

0.021

5. 9x 10- 4

0.018

0.87

0.06

0.87

0.06

6. 3x 10- 4

o. 35

6.8xlo- 4

o. 30

105

42

97

52

*Calculated from equation 4-4.
**To the water column production rate of 210 pb is added an
atmosphe~ic delivery of 15 atoms/cm 2min at site E and of 25
atoms/em min at site S.

199

figures 4-18 and 4-19), the 210pO acti vi tl, is significantly lower
than that of 21 °Pb.
About 50% of the 21 Po appears to have been
lost to raining particles.
The removal time for polonium is thus
about 0.4 years (i. e., similar to that for thorium and significantly less than that for lead).

THE DISTRIBUTION OF RADIUM-226
There is one more member of the uranium series which merits
discussion.
It is 226 Ra , the 1620-year half-life daughter productof 230 Th • Among the constituents of the sea, radium is unique inthat its input is almost entirely from sediments.
Most of Bhe
22 oRa present in the sea is the product of the decay of the 23 Th
which was produced by the de cay of 23 4U dissol ved in the sea andcarried to the sediment on particulates.
As radium is far moresoluble than thorium, much of the 226 Ra "born" in the upper sediments becomes dissolved in the pore water and diffuses back into
the overl,ying water column.
From the average activity ratio of.
226 Ra to 234 U in the sea , it can be shown that about 9 out of
every 100 226 Ra atoms produced from the excess 230 Th in sediments
escapes to the overlying water column.
Because of its half-life of 1620 years, geochemists initiallyhoped that the distribution of 220 Ra might be used as is 14C. todetermine the rate of ventilation of the deep sea (112).
This
hope never came to fruition.
There are several aspects of thegeochemistry of 226 Ra which conspire against its use as a time
tracer.
First, the fact that 226 Ra is added from the sea bottom
reduces its effectiveness relative to 14C~ by roughly the ratio ofthe volume of deep water to the volume of surface water (i. e.,
about 10).
Second, the fact that it is fixed into particles by
surface-dwelling organisms necessitates a large correction forparticulate transport.
Third, its release from sediments shows
large geographic variations which are as yet only qualitatively
understood.
Radium has no stable isotope which might be used as is 12c~ toseparate the influence of particulate transport from the influenceof radioactive decay on the distribution of 14 C . However, as theelements Ra and Ba have nearly identical chemistries, we canconsider 226 Ra to be an "isotope" of barium.
As 14C/ C ratios are
more useful time clocks than absolute 14 C concentrations, so also
should 226 Ra/ Ba ratios be more useful time clocks than absolute
220 Ra concentrations.
In figure 4-20 is shown a plot of 226 Ra versus Ba concentrations for water masses throughout the ocean.
The points withthe lowest 226 Ra and Ba contents are for warm surface waters.
Those with the highest values are for bottom waters in the northeastern Pacific.
The reference line has a slope equal to the
226 Ra / Ba ratio in surface waters.
Were particulate transport the
only process perturbing the distributions of 220 Ra and Ba in thesea, then the points should all lie along this line.
They donot.
Most deep waters have more 226 Ra than expected from theirbarium content.
These anomalies are particularly large in thecase of deep and bottom waters from the eastern Pacific.
Anyapplications of 226 Ra as a tracer of ocean processes must be based

200

A210

-20

A210pO - A210Pb (dpm/IOOkg)-20

O.Or.-r---- •• 0.1-100••- . - . - .

•••••

'

•
I.•• •...•• j --

• ••

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J•

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,

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Pb

(dpm/IOOkg)-10O'-.-.--r.-.
10

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••
•

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pO

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0

10

I

,

~.

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• •••
•
• •
•
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• • , • eIIe
I

4

0.31-

• •• •• .1 ••
0.41•••
N.W. PACIFIC
•
I

-

••

Figure 4-18. Difference between the activi ty of 210 pb and 210 po as a function of
depth in the north tempe ra te Atlantic Ocean.
As can be seen surface wa ters show adeficiency in 210 po activity relative to that of its parent 210 pb •These results
were obtained by Bacon and his colleagues at Woods Hole Oceanographic Insti tution( 149) .
:Figure 11-19. Difference between the activities of 210 Pb and its daughter 210 po in
the upper water column of the north temperate Pacific.
These results were obtainedby the Yale (158) and Hakodate Japan (152) groups.

50• WARM SURFACE WATER• BERING SEA SURFACE WATER• ANTARCTIC SURFACE WATER
o N.W. ATLANTIC DEEP WATER
G S.W. ATLANTIC DEEP WATER
VI N.E. ATLANTIC DEEP WATER
A S.E. ATLANTIC DEEP WATER

40
Ol.::t::.

E

•

$

$

* ANTARCTIC DEEP WATER

--..

0
0
"'-

NORTHEAST
PACIFICBOTTOM WATER$

N.E. PACIFIC DEEP WATER
8 W. EQU. PACI F IC DEEP WATER• E. EQU. PAC IFIC DEEP WATER

$

$

*$

,

88

30
'8

0-

.:g

8

c

0:

$

$

•

U>

N
N

20

.195 dpm/lOOkg
IO-9mol/kg

10~ATLANTIC WARM SURFACE WATER~ PACIFIC WARM SURFACE WATER

30

50

70

90

110

130

150

Sa (lO-9mol/kg)

Figure 4-20. Activi ty of dissolved 226 Ra versus concentration of
barium in waters from throughout the world ocean, as determined on
samples co~lected as part of the GEOSECS program.
The solid line
represents the ra tio of 22 6Ra/Ba in surface wa ters (and hence inparticles formed in surface waters).
The barium results were obtained in the laboratories of L. Chan at LSD and of John Edmond atlVIIT (375).
The 226 Ra analyses for Atlantic Ocean GEOSECS samples
were carried out at the Lamont-Doherty Geological Observatory(116).
Those for the Pacific Ocean GEOSECS samples were carriedou t partly by Chung at the Sc ripps Ins ti tu tion of Oc eanography(117, 118, 160) and partly by Ku at the University of Southern
California (114, 119).

202

on these deviations from the ref~r~nce line. Were there no deviations, then the distribution of 2 Ra would tell us nothing beyond
what we have already learned from the distribution of barium.
As we shall show in the next chapter, the radioactive decayof 226 Ra cannot explain the large departures from the referenceline seen in figure 4-20.
Rather they must be caused by a combination of sluggish interchange between various Rarts of the deepsea and sizable inhomogeneities in the flux of 226 Ra from sediments into the sea.
In particular, the deep waters in the eastern
Pacific must receiv~ a disproportionately large amount of 226Ra .
The flux of 22 Ra is not expected to be uniform over the sea
floor.
The flux of 226 Ra must depend on sedimentation rate.
Thehigher the sedimentation rate, the lower the concentration ofwater column deri ved 23 °Th in the sediment.
The lower the 23 °Th
concentration in the sediment, the lower the concentration of
226 Ra in the sediment pore waters and the lower the diffusive fluxof 226 Ra atoms into the overlying bottom water.
Herein must lie
part of the explanation.
Due to its great breadth, the Pacific
Ocean receives less continental detritus (per unit area) than does
the Atlantic Ocean.
Also, due to its shallower lysocline, lesscalci te accumulates in the Pacific Ocean than in the Atlantic
Ocean. Consequently sedimentation rates in the Pacific are on the
average about one-fifth those in the Atlantic.
The flux of 22 6Ra from sediments will also depend on thefraction of the 226Ra produced by 230Th decay to be released tothe pore water (as opposed to bound in or absorbed onto mineral
grains) .
We have very li ttle knowledge about the dependence of
this fraction on sediment type, and hence we cannot predict theregional variability in this factor.
Fortunately, a means of directly measuring the flux of 226Ra
a toms from sediments exis ts.
It invol ves measurements of the activities of both 226 Ra and 230 Th as a function of depth in marine
sediments.
Normally the upper 10 or so centimeters of sedimentsh~w a deficiency of 226 Ra activity.
The integrated deficiency of
22 Ra activity is a measure of the amount of 226 Ra lost from thesediment. Qochran at Yale (121) and Kadko at Lamont-Doherty (120)have made such measurements.
Large differences from place toplace on the sea floor are found (see table 4-8). Although enoughcores have not been measured to permi t firm conclusions to be
drawn, the cores from the eastern Pacific yield fluxes considerably higher than the global average necessary to maintain the
226Ra inventory in the sea.
The Atlantic cores show values lower
than this mean.
The next step in research in this area must be tobetter document these patterns of 226 Ra input from marine sediments.
Only then will the 226 Ra anomalies find a meaningful
application as an ocean tracer.

ANTHROPOGENIC PLUTONIUM IN THE SEA
As the result of nuclear tests, substantial amounts of theisotopes 239pu and 240 pU were released to the atmosphere during
the last 25 years.
Like the 210 pb atoms generated in the atmosphere by the decay of radon, this plutonium is carried by rainfall
and by aerosol impact to the earth's surface.
Thus a sizable203

Table 4-8. Estimates of the 226 Ra flux from the sea floor.
.
d eep sea
based on measurements of the 226 Ra d e f·lClencylnsediment cores. Results obtained mainly by Cochran while at
Yale University (121).

Ocean

La t.

Long.

Sediment
Type

Sed. Ratecm
10 3 yrs

Flux
atoms
cm 2 minAtl

40 0 N

33°W

CaC0 3 ooze

2.9

5

Atl

29° 8

35°W

CaCa 3 ooze

"'0.1

30

Ind

26° S

74° E

CaCa 3 ooze

0.3

60

Ind

14° 8

78° E

3i0 2 ooze

0.3

80

Ind

39°8

120 0 ERed clay

0.08

60

Ant

53°8

35° E

8i0 2 ooze

1.6

26

Ant

67°8

30° E

Red clay

0.3

45

Pac

9°N

151 ° W

810 2 ooze
+ clay

0.15

190

Pac**

12° N

137° W

8102 ooz e+ clay

"'0.1

420

Pac

lIoN

139°W

Si0 2 ooze
+ clay

"'0.1

400

Pac

15° N

126° W

3102 ooze+ clay

"'0.1

500
80

World Average**Average f~~x of 226 Ra from sediments necessary to balance thedecay of 2 Ra within the sea.
**Result by Kadko of the Lamont-Doherty Geological Observatory(120) •

204

fraction of this plutonium now resides in the ocean.
As shown infigure 4-21, a survey of the Pacific Ocean made by Vaughan Bowenand his associates at Woods Hole Oceanographic Institution as partof the GEOSECS program reveals that as ofplutonium re~i~ed
largely at intermediate depths (164).
Like
8 Th , 239 pu and
4 Pumust be removed onto particles in surface water and then at leastin part released from these particles as they pass through themain thermocline.
Again it seems reasonable to invoke uptake byplants followed by release from falling fecal rna terial for this
distribution.
The sediment trap results obtained in the Pacific (see figure4-22) show that not all the plutonium-bearing particles are destroyed in the thermocline (134). The fact that more plutonium wascaught in the three t raps deployed in deep water than in the two
deployed in the thermocline clearly demonstrates that some of theparticles (fecal pellets) fall straight through to the bottom.

1.J14

TOWARD A MODEL OF METAL TRANSPORTLet us now sit back and try to generalize these findings.
Our first lesson has to do wi th the implications of reversible
absorption.
This is best done by considering an aquarium filledwith sea water.
A pump keeps the water well mixed.
Water thatflows into the aquarium carri~s particles and tracer metals. This
inflow jus t balances evaporation so the volume doesn't change.
The particles settle to the bottom carrying an equilibrium amountof the tracer metals.
Once on the bottom the tracers are assllined
to remain permanently bound on the particles.
The system isallowed to run until steady state is reached (i.e., the metal
tracers leave at the same rate they are added).
Let us see what
determines the residence time of a gi ven metal tracer in thisaquarium. The residence time, T, is given by:
T

4-12

=

where CB and Ci are respectively the concentrations of the
tracer In the aquarium and in the inflow, V is the volume of theaquarium and I is the inflow rate of water.
If the amount of a given metal leaving the tank is to equal
the amount. entering, then:C l"

I - Cp S

4-13

where Cp is the concentration of the metal on the particles and S
is the flux of particles into the tank.
If equilibrium is achieved between the particles and thewater, then:Cp = KC a

4-14

where K is the distribution coefficient of the metal between waterand particles.

205

E

w

LONGITUDE

140

160

180

160

O.O.--..-......- .......-...-....-.....-~ .....- . - . p -.......-E

E-W SECTION
PACIFIC OCEAN

~

~ 0.5

------0.2,-----.-1-------0.4-----_-J

0.4'" 32°N

I0...
W

0.2

239pu a 240 puo...

dpm/IOOkg

1.0
0.0. - - - , , - - -.....- - -.....----t...--...-.IIIIIl~-...
N-S SECTION
PACIFIC OCEAN
I'V:c 0.5
.....
CL
W

ISOoW

'\

1.0

o

239 pu a 240 pu
dpm/IOO kg

1.0

·s

20

o

20
LATITUDE

40

60
N

Fi~ure 4-21. Sections of the concentration of dissolved 239 pu plus

24 Pu across the North Pacific (upper) and along the international
dateline in the North Pacific (lower).
The contours are in
dpm/100 kg.
The concentration of 239pu and 240pu are summed
because the a spectrometric technique used for the measurements
does not resolve these two isotopes.
These measurements were madeby Bowen and his colleagues at Woods Hole Oceanographic (109).

206

239.240 pU (dpm/g)
1.0

O

3.0

2.0

O ~---....------,r------'
I
I
)(

I~

..---.

2

E

-

)(

-

-

~

.......,
I

J-

)(

3

~

-

4

-

-

a..

w

0

x

-

SITE Px

6...-.

. ...-.
1

1......

...

Figure 4-22. Concentrations of 239 pu and 240 pu on material caught
in sediment traps deployed at various water depths at site P at
20 0 N and 150 0 W in the North Pacific.
These measurements were madeby Anderson at Woods Hole Oceanographic (134).

207

Combining these equations, we get:
V

T

4-15

= KS

Thus the residence time is inversely proportional to the particle
flux and to the distribution coefficient.
The inverse dependence
on the particle flux is certainly in accord with the evidence given above.
The dis tribution coefficient dependence could explain
the difference in residence time from one metal to another.
Weshall have more to say abou t the dependence of K on the chemical
properties of metals.
Our second lesson is based on the observations that the fluxof particles shows strong lateral gradients and that the time formixing along isopycnal surfaces is of the same magni tude as thetime for removal of the reactive metals we have been discussing.
To understand the implications of these observations we will consider a model consisting of two coupled aquaria.
As shown in figure 4-23, water is pumped back and forth between these aquaria.
Half of the particle flux is added to the smaller (i.e., coastal)aquarium and half to the larger (i.e., open ocean) aquarium.
The
metals are added in proportion to the volumes of the reservoirs(i.e., 10% to the coastal and 90% to the open ocean aquarium).
The equations governing the metal concentrations in the tworeservoirs are as follows.
First for steady state throughput of
the metals:

4-16
and for steady state in the open ocean reservoir:
0.9M + TC c = 0.5 SKC o + TC o

4-17

where M is the total flux of a gi ven metal to the aquaria, S is
the total particle flux, T is the water transfer rate, K is the
distribution coefficient, and Cc and Co are the steady state
concentra tions of the metal in the coas tal and open ocean aquarium, respectively.
These equations can be solved simultaneously,
yielding:
Cc

M

SK

[ 2 -

o. 9SK + 2T

O.5SK + 2T

J

4-18

and
C o -- MSK[ o. 9SK + 2T
0.5SK + 2T

J

4-19

If the area of the bottom of the coastal reservoir is one tenth
the total area and that of the open ocean reservoir is nine tenths
the total area then the ratio of the rain rate per unit area onto
the floor of the coastal aquarium, Rc , to the rain rate per unit
area onto the floor of the open ocean aquarium, Ro , is as

208

1-5
2

__ T
O.9V

O.IV

T

COASTAL

O.fA,

OPEN OCEAN

O.9A

Figure 4-23. Two-box model designed to show the interac tion between lateral mixing and a geographically non-uniform particle
flux.

209

follows:
2
Rc
Ro

=

0.5 SKCc/A c
0.5 SKCo/A o-9T
0.9 + 2 SKT
0.5 + 2 SK
T
2
0.9 +
SK
T
0.5 + 2 -SK

4-20

The rain rate ratio for various combinations of lateral mixing
rate (i.e., T) and particle rain rate (i.e., S) is as follows:

o

0.3

1

4.2

1

6.5

3
8

00

9

Further if we compare the results for two metals which have K
values differing by a factor of three, one yielding a T/SK of 0.3
and the other a T/SK of 1, the rain rate ratio of the two isotopes(i. e., Rc/R o ) would differ by a factor of 1.5.
The metal wi th
the higher K would have more similar rain rates in the tworeservoirs than the one with the lower K.
Thus we see ~~oa more~uantitative way
how the lateral separation between
Th and31 pa observed in the ocean might be accomplished.
Our third lesson has to do with the use of the distributionsof short-lived tracers to tell us about recycling from the sea
floor.
The point to be made is that these distributions are
misleading in this regard.
To do this let us consider a large,
cone-shaped tank filled wi th sea water.
A stirring device keeps
the water well mixed.
Particles are metered into the water at a constant rate.
These particles settle to the bottom and become part of the sediment.
They are allowed to accumula te until the plug of sedimentat the apex of the cone reaches a certain mass.
At this point
devices are installed which mix the sediment and which remove a
fixed fraction of the sediment each hour.
The fraction is set sothat the amount removed just matches the rain of new material.
At
this point a metering device is turned on which adds three isotopes of the same element to the surface of the wate'r.
Two isotopes are radioacti ve wi th differing half-Ii ves.
One is stable.
The system" is then allowed to run until the tracer distributions
come to steady state.
The operational characteristics of thesystem and the half-Ii ves of the radiotracers are given in table
4-9.
Let us consider the situation where scavenging is irreverssible and the removal probability of the tracers from the watercolumn is 0.10 per hour.
In this case the amount of a given isotope in the water, MW, would be given by:210

Table 4-9. Operational characteristics of the three-isotope tracer

experiment described in the text.
1.00 Cubic meters

Volume1. 00 grams

Steady State
Sediment Load

1.00x10- 3 grams/hr

Sediment Addition
Rate
T Sediment1000 hrs

T1/2 Sediment693 hrs

T Radioisotope A

10 hrsT1/2 Radioisotope A

6.93 hrs

T Radioisotope B

100 hrs

T1/2 Radioisotope B

69.3 hrs.

Table 4-10. Summary of results from the example designed to demons trate the insensi ti vi ty of short-Ii ved tracers as indicators ofrecycling.

Short-lived
IsotopeLong-lived
Isotope

StableIsotope

In Water
No Recycling
Recycling

5.0 R
5.2 R

10 R
110 R

9 R
16 R
In Sediment Mixed Layer

No Recycling
Recycling

5.0 R
4.7 R

82 R
76 R

1000 R
1000 R

Total
No Recycling
Recycling

10 R
9.9 R

91 R
92 R

211

1010 R
1110 RMAW

1
R 0.10+0.10

5R

4-21

MB W

R 0.10+0.01

1

9R

4-22

I

lOR

4-23

MeW = R 0.10+0.00

where R is the rate at which each metal is added to the system andA, B, e refer respecti vely to the short-Ii ved, long-Ii ved andstable isotope.
The left hand number in the denominator is thes ca venging probabili ty; the right hand number is the radiodecay
probability.
The steady state amount of a gi ven isotope in the sediment
MS is given by:

1

MAS = 0.5R 0.001+0.100

5R

MB S

1
0.9R 0.001+0.010

82R

Me S

1
1.0R 0.001+0.000

1000R

4-24

4-25

4-26

The left hand number in the denominator is the sediment removal
probability; the right hand numbe~ is the decay probability.
The
number preceding R is the fraction of the isotope ~eaching thesediment(as opposed to undergoing radiodecay while in thewater).
Now let us consider how things change if recycling occurs.
We will assume that one-hundredth of the sediment burden of theisotope is recycled back to the water each hou~. In this case, at
steady state the following mass balance equations apply.
For thewater:

R + O.OIMS = (0.10 + A)MW

4-27

where the terms on the left give, respectively, the amount of newmaterial added to the water each hour and the amount recycled fromthe sediment each hour, and the term on the right gives the amountof the isotope removed from the water by scavenging and radiodecay
each hour.
For the sediment the steady state balance is given by:

0.1 MW=(0.001 + 0.01 + A) MS

212

4-28

where the term on the left is the addi tion rate from the water~
and the term on the right is the loss rate via sediment removal,
recycling and radiodecay.
Simultaneous solution of these equations yields:
MAW

MB W
and

Me.W

5.2R
16R
110R

4.7R
MaS = 76R
Me S = 1000R
IvIA S

A comparison of these results with those from the simple
scavenging case (see table 4-10) reveals that for the short-lived
isotope the contribution of recycling is only 4% of the total
input.
For the long-lived isotope the contribution of recycling
constitutes about 50% of the total input. For the stable isotope
the recycling constitutes 91% of the total input.
Because of the manner in which the sediment at the base of
the cone was handled (i. e. ~ it was stirred and also graduallyreplaced), one might ask whether this example has anything to dowith the real ocean.
As we shall see in the next chapter, real
sediments are indeed much like our example.
Burrowing organisms
keep the upper 5 to 10 cm of the sediment well mixed.
As newsediment is added, material gets buried beyond the reach of theseburrowers. Hence, as in our cone, there is a reservoir of stirredmaterial from which recycling may occur.
The residence time ofmaterial in this bioturbated layer ranges from about 3500 years
for carbonate oozes to about 20,000 years for red clays.
Any radiotracer wi th a half-life much smaller than the replacement time of the material in the bioturbated layer at the topof the sediment will not provide valid information regarding therecyling of its long-Ii ved or stable counterparts.
210 pb provides
a good example.
Wi th a half-life of only 22 years it will act
much like the short-lived tracer in our example. Its distribution
in the sea has nothing to tell us about the extent to which stable
lead is recycled from the sea floor.
On the other hand, the distributions of 230 Th and 231 pa , which have mean lives greater than
the residence time of material in the sediment mixed layer, do
carry information about recycling.
Our fourth lesson is based on the observation that thedis t ri bu tians of the two ai rborne trace rs, 2 10 Pb and 2 3"~J~ 21+ °pu,
show a prominent difference.
The concentration of 210 Pb isslightly higher at the surface than at 400 meters, while that of239~240pU is higher at 400 meters than at the surface.
What isthis telling us about the difference in behavior of lead andplutonium?
To understand this, let us consider the two-reservoir system
shown in figure 4-24.
Radioisotopes are added to the upper reservoir at the rate M.
C.onsistent with the findings from the distribution of bomb-produced tritium, we allow the waters in theupper reservoir to mix with those in the lower reservoir.
The
metals added to the upper reservoir are assumed to be in part213

M
SURFACE MIXED LAYER

T

T

THERMOCL I NE

M
Figure lJ-~4. Two-box model designed to show the interaction between vertical mixing and a step-wise downward tranport of metals
on particulate debris.

214

extracted from the water by plants and transferred on particles tothe lower reservoir, where the tracers are once again released to
the water.
In the lower reservoir they become attached to raining
particles and are gradually carried to the deep sea.
Neglecting radiodecay, the mass balance of the tracer in theupper reservoir is given by:

4-29
where M is the metal addition rate, T is the water transfer rate,
fU is the fraction of the metal incorporated in plant matter ina unit time, Vu is the volume of the mixed layer and Cu and
CL are respectively the concentrations of the metals in theupper and lower reservoirs.
The mass balance of the metal in the lower reservoir is given
by:

4-30
where fL is the fraction, of the material in the lower reservoir
carried away on particulate matter per unit time, and VL is thevolume of the lower reservoir.
Based on the simultaneous solution of these two equations,the ratio of the concentration of the metal in the waters of thelower reservoir to that in the upper reservoir can be shown to be:

4-31
ThUS,
if T is much larger than fLVL and fUVU'
then theconcentrations in the two reservoirs will be equal.
210 Pb comesclose to this si tua tion.
If T is small compared to f uVu andfLVL then the concentration ratio will depend on the ratio of
these
two
terms.
For
exampleif
fUVU
isgreater
than
fLVL' then the concentration of the metal in the thermocline
reservoir will be higher than that in the mixed layer.
This isthe situation for plutonium. Thus it appears that plutonium has aresidence time in the oceanic mixed layer with respect to removal
on particles which is much smaller than that for lead.
There is, of course, an important difference between thesituations for 239,240 pu and 210 pb .
The former was added over ashort period of time about a decade before the GEOSECB survey was
made.
The latter has been added continuously.
Thus it might beargued that the difference in the vertical distribution of these
two tracers reflects their time histories rather than their particulate affinities.
While this may be the case, in our estimamation the distribution difference is too large to be accounted
for by the time histories alone.
DISTRIB UTIONS OF STABLE METALS IN THE SEA

Let us now consider the distributions of stable trace metals
in the sea.
Thanks to the superb work of Bruland and his co-215

Zn (I0-9mOI/kg)

Cd (lO-IOmol/kg)

10
0
2
4
6
8
\0 0
2
4
6
8
0 Q---.---..---.....-oop--., ~=0I==::r:=--.--.-.......,-,

-2
-....
E

.::.::

J:

~
w 3

0

4
5
Ni OO-9mol/kg)

0
0

4

6

8

Cu OO-9mollkg)

10

2

:3

4

5

6

-~ 2
....J:

a. 3
w

o

4

Figure 4-25. Plots of dissolved Zn, Cd, Ni and Cu as a function of
depth in the central North Pacific.
These resul ts were obtainedby Bruland and his coworkers at the Universi ty of California at
Santa Cruz (455).

216

I 2 r--...,.-......-..,...-.,....--.,~--r"-..,

C'

8

I

oX

~
0

E

0

<:)

<:)

6

'0

<:>
"'0

~

U

N

4

2

160
40
80
120
H 4 Si04 (I0-6 mollkg )

I

2

3

P04 (I0-6mollkg)

14
40

12
_10

~30
.......

<:)C'.:.:::

<:)

"

0

0

<n

E

0

lD

E

I

0
rtl

Z

0

z

4

10

2

10
5
Ba (I0-8mollkg)

o ..

15

0

I
2
P0 4 (I0-6mol/kg)

3

Figure 4-26. Plots of the data shown in figure 4-25 against thatfor nutrient elements showing similarities in their depth distribution.
For the pairs Zn-H 4Si0 4 and Cd-P0 4 the covariance isalmost perfect (better, for example, than even that for N03-P04
shown at the lower right). By contrast the plot for Ni versus Ba
i s quite nonlinear.
Thus, while Ni and Ba are both biointermediate elements showing about the same removal fraction in surface
waters, their release patterns from particles cannot be the same.

217

workers at the University of California at Santa Cruz, there are
four metals for which excellent data are available (455).
Asshown in figure 4-25, the vertical distributions of Zn, Cd, Ni and
Cu in the North Pacific Ocean are similar to those for the nutrient constituents.
In fact, that for Cd has precisely the shape ofthat for P0 4 (see figure 4-26), and that for Zn has precisely thesame shape as that for H4 Si0 4 (see figure 4-26).
These resemblances require that both Cd and Zn be totally removed onto particulate matter in the surface ocean and that they have regeneration
functions identical to those for P0 4 and H 4Si0 4 . As shown in figure 4-26, Bruland and his coworkers demonstrated that the correlation between C~d and PO 4 is even better than that between NO 3 andP0 4 •
While there is no simple biological or chemical explanation
for these findings, they do tell us that the formation and destruction of living matter must be very important in determining
the pathways followed by these elements through the sea.
Nickel is a biointe rmedia te element.
Its concent ration inwarm surface is about 20% that in deep Pacific water.
In thisregard it is similar to barium. However, as shown in figure 4-26,
the two distributions do not correlate as do those of C.d and PO 4
and of Zn and H4 Si0 4 • As can be seen in figure 4-25, the shape ofnickel versus depth curve is similar to that of Cd in the upper
1000 meters and that of Zn in deep water.
Copper (figure 4-25) shows a rather regular increase fromsurface to deep water.
Its pattern is not a replica of that ofany of the nutrient consti tuents.
The copper content of surface
water in the North Pacific averages about 10% that in bottom
water.
Hence, using the definitions given in the first chapter,C~u isalso a biointermediate element.
The shape of the copper
profile has a characteristic not found for any nutrient.
Whenplotted against potential temperature, the curve has a concave
shape (see figure 4-27).
The copper content of the water for any
intermediate temperature is less than would be expected from mixing the overlying and underlying water.
By contrast, the otherconstituents whose profiles we have examined all show a convex
shape when plotted against potential temperature (i.e.,thecontents of these constituents is greater at any given depth thanwould be expected from the mixing of overlying and underlying
waters) .
Thus there must be a net removal of copper from theinterior of the deep water column and a net addition of the othercomponents.
Patterson and Schaule of Cal Tech (436) have carried out avery diffi.cult set of measurements on a fifth metal, lead.
Asshown in figure 4-28, its vertical distribution in the North
Pacific is very different from that of four other elements studied
by B ruland and his coworkers.
Its high concentration in the upper
waters must be anthropogenic in origin.
It comes mainly from theanti-knock lead additive which, until a few years ago, was used inall gasolines.
Stable lead is distributed with depth in a similarmanner to anthropogenic tritium!
The analogy between the distributions of Cd and P0 4 and between Zn and H4Si04 strongly suggest that these metals are recycled many times wi thin the sea as are the nutrient elements.
Were this not the case, then the signatures of the inputs of these218

Cu (I0-9mol/kg)
20".........

U

3

4

5

10

0

"'-'

W
0::
::J

8

~

0::

W

a..

•

6

~

w
t- 4
.-J

«

rZ

2

w

ra.. 00

Figure 4-27. Plot of the Bruland et al. dissolved copper concentration da ta 'as a func tion of potential tempera ture in the central
North Pacific.
The concave shape requires that a net scavenging
of copper occurs in the deep water column.

219

E

~

2

..........

I
I-

a..
w
o

:3

4

5

Figure 4-28. Distribution of dissolved lead wit~ depth in thenorth temperate Pacific Ocean.
The excess for the upper waters isthought to represent pollution (i.e., lead from gasoline).
These
measurements were made by Patterson and Schaule at California Institute of Technology (436).

220

elements to the sea would disturb the correlations.
In thisregard Zn and C:d appear to differ greatly from Th and Pb, whichgive the appearance of being removed rapidly and irreversibly fromthe sea.
This basic difference is confirmed by the estimates ofriver input given in table 4-11. Although the river data (obtained by Ed Boyle and his coworkers at MIT, 632) is for the Amazon
only, there is no question that Cd, Cu and Ni reside in the sea
longer than the ocean's mixing time.
One last point should be made about these five metals before
going on.
They all show unexpectedly high concentrations ina tmospheric dust and aerosols.
In table 4-12 the ratio of theconcentrations of these fi ve metals to aluminum in material filtered from the air and in residues from polar ice is compared totheir ratio in average crustal material.
In all cases these fiveelements are greatly enriched.
While the enrichment for lead in
the air samples and in the ice sample from Antarctica is likely
the result of contamination with gasoline lead, that in the pre1900 ice sample from Greenland suggests that lead also shows anatural enrichment.
Selenium shows an even stronger enrichment
than do the five stable metals discussed here.
While the originof these enrichments is not known, the suspicion is that they are
related to the biological affini ties of these elements.
The release to the air could be either from terrestrial vegetation orfrom froth blown off the sea surface.
As can be seen from thetable, not all metals show this enrichment.
Some have ratios toaluminum close to the crustal average.

STABLE ISOTOPE RATIOS IN REACTIVE METALSIsotopic separations similar to those observed for oxygen,hydrogen and carbon are not observed for the metals of interest to
us in this chapter.
The absence of such separations is in partdue to the fact that the magnitude of these separations depends onthe fractional mass difference between isotopes (i.e., on ~m/m).
For the hydrogen-deuterium pair this difference is very large( i . e • , ( 2-1 )/ 1 ) • For the 1 3C. - 1 2C Pair i t i s about o . 08 ( i • e . ,( 13 -12 ) / 1 2 ) an d for the 1 8 0 - 1 6 0 pa i r i t i s abo u t o . 1 2 ( i • e. ,(18-16)/16).
By contrast for metals with masses in the range 56(iron) to 207 (lead) the fractional mass difference rarely exceeds
0.03.
Further, the chemical transformation which these metals
undergo in nature do not produce as large fractionations for agiven mass fraction change as do some of the reactions experienced
by carbon,- hydrogen and oxygen.
In any case, no one has found any
measurable variations in isotopic composition which would prove
useful in marine chemistry.
There are two exceptions, the elements lead and neodymium.
Both show large natural variations in isotopic composition.
These
variations are, however, not the result of isotope separations.
Rather they result from the fact that three of the four lead isotopes and one of the neodymium isotopes has been in part produced
within the earth by the
decay of long-Ii ved radiois otopes.
The
isotopic compositions vary because of chemical separations between
the parent and daughter isotopes within the earth.
In the case of
Pb the parents are isotopes of the elements U and Th.
Planetary221

Table 4-11. Residence time estimates for the elements dealt
with in this chapter.

Element

Conc. +
Rivers10- 9 moles/kg

Conc. ++
Sea
10- 9 moles/kg

River
Input
moles/yr

Mass
In Sea
moles

-

-

-

24.0
<0.1
6.0

3.0
0.6
7.0

8x10 8<0. 3x 10 7
2x10 8

Th
Pb
Cu
Cd
Ni

-

1"

yr
50*
40* *
5000
>3x 10 5
5xlO 4-4x 10 12
Ix 10 129x 10 12

*Based on 23 °Th da ta.
**Based on 210Pb data. Because the half-life of 210Pb is only 22years, its distribution does not provide a measure of the importance of recycling.
Hence, the 1" for stable lead may be larger
than 40 years.
Indeed the value for stable lead in table 1-6 is1100 years.
~ased on Amazon River data obtained by Boyle and his coworkers at
MIT (632).
+~ ased on the measurements of B ruland at the Uni versi ty of C:alifornia Santa Cruz (455).
Table 4-12. Enrichment factors (relative to the mean crustal
abundance) of various elements in material filtered from theatmosphere and in the atmospheric fallout onto the polar ice
caps.
In each case the concentra tions have been normaliz ed tothat of aluminum* as summarized by Li (652).

*

Element

N. Temp.
Aerosol

Greenland
Ice(Pre-1900)

Antarctic
Ice

Thorium
Iron
Manganese
Vanadium
Samarium
BariumNickelCopperZinc
Cadmium
Lead
Selenium

1.2
1.6
0.5
1.6
2.0

1.2
1.0
1.0
1.0

1.0
1.2
1.2
1.0

.

-

25
30
300

-

1200**
1000

-

-

1.0
25
16
10
200
20
1000

2.0
50
40
6000
200**
500

_ ([Element]/[AI]) sample

Enrichment factor -([ElementJ/[Al])crust**Anthropogenic contamination.

222

differentiation has led to a large enrichment of U and Th in the
grani tic crust and presumably to an enrichment of lead in thecore.
Thus lead which has spent a subs tantial amount of time inthe earth's grani tic crus t, has high ratios of 20 6 pb , 20 7 Pb and
208 Pb (the radiogenic isotopes) to 204Pb (the non-radiogenic
isotope) .
By contrast that lead which has recently reached thecrust from the mantle has lower ratios.
Planetary differentiation
has enriched both samarium (the parent element of the radiogenic
isotope 143 Nd ) and neodymium in the earth's crust.
However, thedegree of enrichment is greater for neodymium than for samarium.
Thus neod:t:mium from the old portions of the grani tic crust has alower 14 Nd (radiogenic) to 144 Nd (non-radiogenic) ratio thanneodymium recently derived from the mantle.
Measurements of lead and neodymium isotopes in seawater, inmarine hydrothermal deposits, and in manganese nodules have muchto offer in two areas. First, they offer a means of distinguishing the contributions of the two major natural sources of these
metals, namely, continental weathering and ridge crest hydrotheral acti vi ty.
Neodymium from ridge cres ts is deri ved from the
mantle.
Thus its isotopic compos i tion is qui te different fromthat for continental rocks.
In the case of lead the situation ismore complicated.
Most of the lead dissolved in the upper sea is
now anthropogenic (from the gasoline additive lead tetraethyl).
This contamination has not, however, measurably altered the isotopic composi tion of marine sediments and nodules.
Further, as
most of the gasoline lead is manufactured from sedimentary leadores with an unusual isotopic composition, it might be possible touse isotopic measurements to separate the contributions of thesevarious sources to the lead currently dissolved in the deep sea.
The second application is to use the isotopic composition oflead and neodymium to place limits on the residence time of theseelements in the sea. Because of the reactivity of these elements
it is difficul t to determine how much of the lead and neodymium
carried by rivers will become dissolved in the sea and how much
will follow the detrital material directly to the sediments.
Similarly, it is difficult to determine what fraction of these elements in sediments is "detrital" and what fraction is "labile."
Thus conventional estimates of residence time for these elements
are not very reliable.
Isotopic evidence can be used in the following way in this connection. Because of geological differencesamong the ocean basins, the isotopic composition of the lead andneodymium ~hey receive will differ. For example, the Atlantic issurrounded by old continents and its ridges have slow spreading
rates.
By contrast the Pacific is surrounded by much younger
rocks and its ridges have high spreading rates. Thus the lead andneodymiumentering
theAtlantic
shouldhave
anisotopic
composition more like those for average continental material while
the lead and neodymium entering the Pacific should have an isotopic composition more like that for the mantle.
Now, if in addition
the residence time of these elements in the sea (i. e., the time
f rom entry to final burial) is long compared to the interocean
mixing time (i.e., ~ 500 years), then these differences should besmoothed out.
On the other hand, if the residence time is shortcompared to interocean mixing time, substantial ocean to ocean

223

SEA WATER

A- AA· A·

P•
p. p.

A• A·

I
-15

I

I

-10

-5

I

a

MANGANESE NODULESHYDROTHERMAL CRUSTSMETALL J FEROUS SEDS. p. .p
p•
I·
•p
p. .p
A·
I·

A· A• A·I•
A· A· A· r· eI I·
A· A .A .I .A lANl!

.r

A·

I.

I

I

~NT .p .pI-10-5
E 143 Nd (%00)-15

Ft~urep• .p
p• p. .p

p.
p.
.p
I

o

of the 143 Nd/ 144 Nd data (expressed asE
3 Nd following the definition given in table 4-13) for seawaterand authigenic ferromanganese deposi ts from throughout the world
ocean.
The letter designations represent the ocean from which thesample was taken.
The dots designate the laboratory in which thedata were obtained.
Those marked in the upper right are analyses
made at C~l Tech (333,336,337) and those marked in the lower left
are analyses made at the Lamont-Doherty Geological Observatory

4-29. Histogram(339, 344).

224

11-13. Geochemical abundances of neodymiwn and samarium andthe isotopic composition of neodymium.

Table

143 Nd*10-~~/gNd
10- 6 g / g

Nd
Sm

Average
Earth

0.32

0.97

3

0

Mid-ocean
Basalts

3.5

10

3

+10

Grani tic
Crust

4.5

27

6

-20

Atlantic
Ocean

-

-

-

-11

Pacific
Ocean

-

-

-

- 3

*

143 Nd )144 Nd143 Nd )
144Nd

samp

14 3Nd )
144Nd

225

std

std

E

0/000

differences would be expected.
Let us first examine the isotopic data for neodymium.
Measurements by the Cal Tech group (333,336,337) and by the Lamontgroup (339) on ferromanganese deposits and water samples are summarized in figure 4-29.
These data show clear differences between
the Atlantic and Pacific Oceans and hence suggest that residence
time of neodymium* in the sea is smaller than the interocean
mixing time.
The sense of the differences is consistent with that
expected from the geological argument presented above (see table4-13).
Analogous isotopic patterns would be expected for lead.
As
we have seen, this element is removed from the sea on a time scaleof several
decades.
Hence, if recycling from the sea floor is
not important then the isotopic composition of the lead in manganese nodules should be similar to that of the mixture of leadreaching a gi ven basin from its surrounding continents and fromits ridge crests.
As shown in figure 4-30, Patterson and Chow at
Cal Tech (341,342) demonstrated this to be the case.
Nodules fromthe Atlantic have lead richer in the radiogenic lead than nodules
from the Pacific.

TRANSPORT OF IRON AND MANGANESE IN THE SEA
One of the moti vations for looking into the transport ofmetals through the sea was to understand the growth of manganese
nodules on the sea floor and the enrichment of Fe and Mn in red
clays.
So far we have not presented any data for either of theseelements.
No da ta exis ts for iron.
The reason is that iron has anextremely low concentration in seawater and that iron contamination on ships and in labs is large and difficult to control.
No
one has yet been able to beat this nearly insuperable combinationof difficulties.
From laboratory and coastal ecosystem tracer studies there isgood reason to believe that iron will act much like thorium in sea
water (i.e., it will within a few decades be removed to the sediment) .
However, unlike thorium, iron may become reduced in thesediments and remobilized.
Hence it is likely seeping into thesea from sediments rich in organic matter as well as entering fromrivers and ridge crests.
The concentration of Mn has been determined in seawater.
The
data in hand strongly suggest that, once in the sea, manganese,like thori.um and lead, is rapidly removed to the sediment.
Asshown in figure 4-31, the concentration of manganese is higher insurface water than in deep water and higher in surface waters near
the continent margin than in open ocean surface waters (449).
Hence the manganese released from coastal sediments and added tothe sea from rivers never becomes thoroughly mixed wi thin thesea.
Rather, it is captured on particulate debris and taken to
the sea floor.
Dramatic evidence for the rapid removal of manganese from the*As manganese nodules contain ten times more neodymium than sediments, this Nd is likely derived from solution.

226

MID OCEAN RIDGEBASALTS(>CONTINENTAL
LEADS

•
en

w

...J
::::::>
Q

o
Z

lL.

o

a:
w
CD
~

~ ATLANTICDPACIFIC::::::>

z

•
•
18.0

18.5

19.0

19.5

206Pb/ 204 Pb

Figure 4-30. Histogram of 206 Pb to 204 Pb ratios in manganese
nodules from the Atlantic and Pacific Oceans.
As uranium is enriched over lead in the earth's continental crust relative to themantle" the ratio of 206 Pb (produced by the decay of 238 U) to
204 Pb will be higher in continental leads than in leads recently
derived from the mantle (i. e., leads in mid-ocean ridge basal ts) .
It can be seen from this histogram that Atlantic leads have alarger continental component than do Pacific leads.
Most of theanalyses shown here were carried out at Cal Tech by Chow andPatterson (341,342).
Thosesquares with a dot in the upper
right hand corner represent analyses carried out by 0 'Nions andhis colleagues at the Lamont-Doherty Geological Observatory (344)
and those wi th a dot in the lower left hand corner analyses byReynolds and Dasch at Oregon State University (343).

227

DISSOLVED Mn (l0-9mol/kg)
0.0

0.5

a

1.0

-E 2
-

E

~

~

~

~

t-

I-

o.
W

0

DISSOLVED Mn (I0-9mollkg)
5
10

a. 3

:3

lJJ
0

NORTH
PACIFIC
33°N 145°W

4

NORTH
PACIFIC
37°N 123°W

4

5

5

5
oa

SURFACE
WATERSNORTH PACIFIC~.......

4

0

E

'"0
I

<:>
(')

]>

@ r

"0:0

:3

z

c

:E
0
UJ

<:>

2

0
0
0

>
-1
0

en
en

0

-0

0

0
0

a

150

% 0 0

0~

0

140
130
LONGI TUDE (OW)

\
o 0

-

]>

0

(')

0

l>

en

-..
120

Figure 4-31. Upper panels show vertical profiles of manganese in
the North Pacific Ocean.
Tha t on the left is for an open oceans ta tion and that on the right for a coas tal s ta tion.
Note theten-fold scale difference between these two diagrams.
The coastal
waters carry an order of magnitude more Mn!
The lower panel shows
the decrease in the Mn content of surface water away from thecoast of California.
The results are the work of Landing and Bruland of the University of California Santa Cruz (449).

228

Mn (p.g/kg)
50

-E

0.2

0.4

0.6

E

-

-

~

~

:c 2
l-

a..

W
0

0.0

• STN 2 21 0 N

EPR CREST

3

o STN 4 23°N
TAMAYO F.Z.

Figure 4-32. Vertical profiles of excess 3He and of dissolved
manganese for a station right over the crest of the East Pacific
Rise at 21 0 N (solid circles).
Injection of new 3He and manganesefrom the rise cres t can be clearly seen.
3 He
resul ts from astation at some distance from the rise crest (open circles)
establish the ambient 3 He background.
The samples were collectedby Weiss of the Scripps Institution of Oceanography (447).
Themanganese measurements were made by Klinkhammer of the University
of Rhode Island (446).
The heliwn isotope measurements were madeunder the supervision of Lupton in the laboratory of Craig at
Scripps Institution of Oceanography.

229

inter'ior of the sea is shown in figure 4-32.
As can be seen,
samples taken above the East Pacific Rise show strong anomalies inboth 3 He and manganese (446). Both substances must hence be added
to the sea locally from hydrothermal vents.
Unlike 3 He , whichshows a broad maximum produced by the integrated imprint of hydr'othermal activity all along the Rise crest, manganese shows no suchfeature.
As shown by Weiss of SCr'ipps, this difference can only
be explained if manganese is rapidly removed from the water column(447) .
LESSONS FROM CONTROLLED ECOSYSTEM STUDIES
Experiments car'ried out in large tanks designed to duplicate
the conditions in Nar'ragansett Bay shed light on the relative absorption tendencies of a number of the elements discussed above.
These experiments, conducted by Santschi and his coworkers at theLamont-Doherty Geological Observatory, involve measurements of theremoval times of y-ray-emi tting tracers to the sediments at thebottom of these tanks (635).
It was found (see table 4-14) thatthe elements Fe, Pb, Po, Th, Pa and Pu were all removed far more
rapidly than the elements Cd, Zn, Se, Ba and Ra.
Mn and Co hadintermediate removal times.
These results are consistent with ouroceanographic findings.

DISTRIBUTION COEFFICIENTSAs we showed above, the absorption of metals onto particulatematter appears to be reversible.
If so, this absorption can becharacterized by distribution coefficients.
Chemists working in
the laboratory, using the oxides of manganese, silica, aluminumand iron as absorption media, have measured the absorption tendency for a variety of metals (684).
They find a correIa tion between the absorption tendency and the first hydrolysis constant of
the metal.
The larger the hydrolysis constant, the more strongly
is a metal absorbed.
Iron has the highest hydrolysis constant of
all the metals discussed here.
That for thorium is one order' ofmagni tude smaller.
Those for other of the elements dealt wi th
her'e are listed in table 4-15.
As can be seen, there is somesimilari ty between the order of these constants and the r'emoval
times of elements from the sea. Manganese, however, is so far outof or'der that one wonder's how applicable these constants are in
representing the r'eactivity of elements in the sea.
Uptake byorganisms,' oxidation-reduction reactions and stabilization of
metals by complexing ligands are possible causes for the lack ofagreement.

SUMMARY
We have seen in this chapter that the flow of metals throughthe sea is complex.
It involves both transport on particles andmovement with the water. Because of differences in chemistry eachmetal has its own set of pathways.
Thorium, protactinium, lead,polonium and plutonium are rapidly removed to the sediments.
r~anganese andiron likely have ocean geochemistries similar to230

Table 11-14. Removal time groupings for the elements studied in
MERL tanks by Santschi and his coworkers from the Lamont-Doherty
Geological Observatory (133).
These tanks are designed to
duplica te the si tua tion in Narragansett Bay.
The resul ts shown
here were obtained by studying the removal to the sediments ofradiotracers added to the waters in these tanks.

Group

Half Removal Time
WinterSummerdays
days

Elements

1

Fe, Pb, Po, Th,
Pa, Pu

10-20

2- 3

2

Mn, Co

30-50

5-30

3

Cd, Zn, Se

60-10010-50

4Ba, Ra150-600

80-150

Table 11-15. Hydrolysis constants forthis chapter.
Element

the elements of interest inFirst Hydrolysis
Constant

Fe

,..

10- 3

Th

...,

10- 4

Pb

,..

10- 7

Nd

,..

10- 8

Co

...,

10- 9

Ni

...,

10- 9

Zn

,..

10- 9

Mn

...,

10- 11231

those for these highly reactive elements.
By contrast, copper,
zinc, cadmium and nickel remain in the sea for many mixing cycles.
Except for copper, whose distribution shows definite evidence forscavenging from the deep water column, these elements appear to be
invol ved mainly wi th the biological cycles of the sea.
While
these studies give us some clues regarding the processes responsible for the enrichment of the elements Fe, Mn, Ni, Cil, Co and Zn
in manganese nodules and red clays, they fall well short of providing the answers we seek.
This young science will have to begreatly expanded before these answers are forthcoming.
PROBLEMS

4-1

Shale contains on the average about 3x10- 6 g U/g.
Asswnethat soils have the same uranium content.
C~nsider two
soils, one dry in the upper 20 em and the other water-saturated to the top.
Both have porosi ties of 50% (i. e., they
are 50% by volume pore space and 50% by volume mineral ma~­ter) .
The mean dens i ty of the mine ral grains is 2.5 gl em .
Assume that the activity of 226 Ra in the soil is the sameas that of 238 U and that one-third of the radon produced in
the soils is released to the pore space.
Calculate the fluxes of radon atoms from the two soils (the molecular diffusi~­
ity of radon in water is 1.4x10- 5 cm 2 /sec and in air 0.1 em I
sec).
Also calculate the characteristic depth of radon depletion in the water saturated soil (i.e., the depth where thedepletion is lie as great as at the top of the soil). Compare
the fluxes obtained wi th the fallout ra tes of 21 °pb atoms
shown in table 4-6.

4-2

Sediment traps are deployed for one month at a depth of 75
meters in a saline desert lake.
Material caught in theset raps yields the following ac ti vi ty ra tios for the uranium
series isotopes:
234Th
230 Th=25*

228 Th230 Th

=3.5xl0231 Pa
230Th
4

210Pb
230 Th

*

0.11

=

0.13

Analyses of the water yield the following results:
U = 108xlO- 6 g/liter
234U/238U activity ratio of unity
226 Ra15 dpm/liter
228 Ra = 3 dpm/liter
Assuming that these concentrations are all at steady state,calculate the apparent residence times of Th, Pa and Pb in
the lake waters.
The 21 Pb fallou t rate over the lake is 50
atoms/cm 2 minute.

°

232

4-3

If the mean 226Ra activity in sea water were 50 dpm/l00
li ters, what would be the mean flux of 226 Ra flux fromthe sea floor necessary to maintain this concentration?
The volume-to-area ratio for the ocean is 3800 meters.
What would the "equili brium" 2 26 Ra content of the sediment pore waters have to be to support this flux (by"equilibrium" is meant that which would be present were
there no loss to the overlying sea). How deep would thezone of 226 Ra depletion in the pore waters extend?
Assume a porosity of 50 percent and a pore water resaturation time of one day.

4-4

If the average sedimentation rate in the ocean is 1.5
g/cm 2
10 3 yrs, the uranium content of sea water is
6
3.2xl0- g/liter and the 234 U/ 238 U activity ratio in sea
water is 1.16, then what must be the concentration ofwater column derived 230 Th in the average newly formed
sediment at a depth of 4000 meters? Give your answer inunits of dpm/g.

4-5

If the excess 210 Pb atoms carried to the sea floor on
particulate matter are homogenized throughout the upper
3 cm of the sediment, then what would be the concentration of these atoms in the core top material from theNorth Pacif ic Ocean?
As sume that there is 1.0 g ofsolid matter per cubic centimeter of sediment.

SUPERPROBLEM#4Intrigued by Theft's four-reservoir aquarium, Spencer Crane
decides to build a similar device to study the uptake of reactive
metals.
He starts with a two-reservoir system with sediment at
the bottom of the lower aquarium.
Following Theft's lead, he puts
mixers in each reservoir to keep the water homogeneous, and also a
device which will keep the sediment well stirred.
Pumps are installed to transfer water (and the suspended matter in this water)
between the two reservoirs.
He fills the reservoirs with waterand then adds HgCl2 to poison any organisms it might contain.
He
then ins talls devices to inject particulate rna tter into the surface waters and filtering devices similar to those used by Theft.
The material recovered by the filter in the upper reservoir isregularly' "back-flushed" into the lower reservoir.
The material
recovered by the filter in the lower reservoir is regularly
scraped off and mixed into the sediment.
Another gadget removes
material from the well mixed sediment (so that the amount of sediment remains constant).
The operating characteristics of this
two-reservoir system are as follows:
Volume upper reservoir
1 m3
9 m3
Volume lower reservoir
Water exchange rate
0.1 m3 /hr= 100 gMass sedimentlxlO- 3 g/h r
Rate of particle injection
Filtering rate in upper reservoir
0.3 m 3 /hr

233

Filtering rate in lower reservoir= 0.3 m 3/hr

To simulate the in situ production of short-lived isotopes to thesea, Crane meters isotopes into both the upper and lower reservoirs.
The addi tion rate to the lower reservoir is nine times
greater than to the upper reservoir (so that the addition rate per
uni t volume is the same in both reservoirs).
For this purpose
C~rane selects a reacti ve metal wi th two radioisotopes of differing
half-life (6.93 hours and 69i hours).
The injection rate of eachisotope is 2.0 x 10 8 atoms/m hr (referred to as I below).
Once everything is set up, Crane switches things on and thenlets his J-machine, as he calls it, run until the distribution of
the isotopes between the three reservoirs (i.e., upper water, lower water and sediment) comes to steady state.
C:rane then determines the concentration of each isotope in the upper and lower
reservoir in particulate form.
For the long-lived isotope, hefinds:
CupperClower

9. 3 I

39.3 I

Crane assumes that the removal onto particles is reversible (i.e.,
the isotope achieves an equilibrium distributiton between theparticles and the water).
He describes this distribution by thefraction, f, of the tracer on the particles (i.e., f = amt on
particles/amt on particles + amt dissolved in the water).
He assumes that this fraction is the same in both reservoirs.
Cranefurther assumes that a certain fraction, g, of the isotope in thesediment is released to the water each hour.
From the results on
the two tracers what are the values of f and g?
What are theconcentrations of the short-lived isotope?
A colleague named Suess Banjo, who is interested in the roleorganisms play in determining the chemical pathways through thesea, looks in on C:rane' s experiment and suggests that he repeat it\tith plants, animals and bacteria present in the system.
Crane
agrees and replaces the poisoned sea water and sediment.
Headjusts the nutrient content of the water till he gets the desired
plant growth rate in the upper reservoir and makes sure that the
animals and bacteria in the lower reservoir consume any organic
debris received during the back-flushing of the filters.
He thenstarts a new tracer experiment. When steady state is established,
he performs the same analysis as before. For the short-lived isotope, he finds the following concentrations in the particulatematter:CupperClower

0.0 I

3.1 I

Crane interprets these results to indicate that removal by plants
overwhelms removal on inorganic particles for both of the isotopes
in the upper reservoir.
He assumes that, when the organic material formed in the upper reservoir is destroyed in the lower reservoir or in the sediment, the isotopes are released (those released234

to the water dissolve and are mixed throughout the lower reservoir
and those released in the sediment become attached to inorganic
particles and are mixed throughout the sediment reservoir).
Hefurther assumes that all the organic material is quickly eaten.
Finally, he assumes that the distribution coefficient for inorganic particles established by his first experiment remains thesame and that the fraction of the isotope recycled from the sediment per unit time established in the first experiment also stays
the same.
What must be the fraction of the metal released by theoxidations of organic residues to the sediment as opposed to that
released to the water in the lower reservoir?

LAST MINUTE ADDITION (SEPTEMBER 1982)
Wasserburg and his colleagues at Cal Tech have demonstrated
that the isotopic composition of Nd in Drake Passage deep waterplaces restrictions on the ratio of the contributions of new waterfrom the northern Atlantic Ocean to recirculated water from thedeep Indian and Pacific Oceans to the deep circumpolar flow (702).
If this approach proves valid it may resolve an important questionof long standing in physical oceanography (see chapter 7).

235

HOW FAST DOES THE MILL GRIND?
RATES OF VERTICAL MIXING AND SEDIMENT ACCUMULATIONINTRODUCTION
In chapter 1, we state that the residence time of water inthe deep sea is about 1000 years.
In chapter 2, we state that therate of rain of GaGa 3 toward the deep sea floor averages about 1
g/cm 2 10 3 yrs.
How were these rates established?
Neither thepattern of composition of sea salt nor of the major sedimentarycomponents tell us anything about the rate at which the systemruns.
As we will see in this chapter, these rates are based primarily on the measurement of trace quantities of a radioisotope ofcarbon which travel through the sea and into the sediments alongwith the nutrient constituents.
Let us first see how the distribution of this isotope can be made to divulge the residence timeof water in the deep sea.
Then we will consider how it can beused to determine the rate at which GaGa 3 bearing the sedimentaccumulates.

RATE OF VERTICAL MIXING
A radioisotope of any given element has a chemistry nearly
identical to that of the element's nonradioacti ve or "normal"
isotope. *So a radioactive carbon atom moves through the ocean
just as any other carbon atom.
Since we have already worked ou t
the cycle of carbon in sea water, we can use this information inconnection with the application of carbon-14 (14 C) as a timekeeper.
Radioactive carbon is continually being produced in ouratmosphere.
There, cosmic rays from interstellar space encounter
atoms and fragment some of them, releasing neutrons from their
nuclei.
Many of these neutrons find their way to nitrogen atoms*Small separations between the isotopes of carbon occur duringchemical reac t ions.
These so-ca lied "isotope frac t ionat ions" areon the order of a few percent.
Although we will have a fair amount to say about these fractionations in tre chapters to come it
suffices here to say that in the case of
4 C they are merely anuisance in that we are interested in differences produced byradiodecay.
Fortunately, the 13 c/ 12 c ratios in the samples of14.
14
carbon used for
C studIes can be used to correct thec/cratios to what they would have been in the absence of any isotope
1eparation.
Theory tells us th t the separation between
4C and2C will be twice that between 11C and 12 C •236

(ni trogen gas makes up 80 percent of our atmosphere).
Upon
entering a nitrogen-14 nucleus this neutron can knock out a proton
and remain behind in its place (like a cueball in pool).
Now, asthe nucleus of a ni trogen a tom has 7 protons and the nucleus of a
C atom has 6 protons, in losing one proton the nitrogen nucleus
changes into a carbon nucleus.
The number of nucleons (that is,
the number of neu trons plus protons) in the new nucleus is still
14, the same number as in the old nucleus.
The difference is that
one neutron has entered the nucleus and one proton has been expelled.
The 14 C nuclei so produced are unstable; the mixture of 8
neutrons and 6 protons is not a durable one and endeavors tochange itself back into the stable 7-7 configuration of 14 N •To
do so, the 14 C nucleus must hurl out an electron.
In so doing it
converts one of its neu trons into a proton and the nucleus again
assumes the form of 14 N• This cycle is shown in figure 5-1.
About 100 14 C a toms are being generated by cosmic rays ove reach square centimeter of the earth's surface each minu teeOvermany thousand of years, the amount of 14 C on the surface of theearth has become nearly cons tanto *14 C is disappearing by its
own radioactive decay at the same rate that it is being producedby cosmic rays (i.e., 100 atoms of 14 C undergo radioactive decay
per square centimeter of earth surface each minute).
The property that makes 14 C a very interesting and important
isotope in natural science investigations is that the average 14 C
a tom requires 8200 years to under~o the transformation back to
14 N!
Thus, in their lifetimes, 1 C atoms have a chance to mix
reasonably well but not completely with the ordinary carbon in theocean-atmosphere system.
The fact that the mixing is not complete
makes 14C a very valuable tool in oceanography.
-Before exploring the tracer potential of natural 14 C , let us
first review the rules governing radioactive decay.
The disappearance of a batch of radioactive atoms is a logarithmic process:a set fraction of the atoms undergo radioactive transformationeach year.
For 14C, this fraction is 1 atom in 8200.
This logari thmic decay pattern gives rise to the concept of half-life.
If
we were to start wi th a given number of radiocarbon atoms today
and let them age for 5700 years, half of them would have become
14 N atoms and the other half would s till be 14 C atoms.
Half ofthe half of 14 C atoms that remained would undergo transformationin the next 5700 years, and so forth.
For an analogous si tua tion, consider a roomful of students.
Each has a coin which he fl ips once eve ry second.
No one can*The concentration of a radioisotope is often expressed in radioactivity rather than mass units. Since the mass of a given radioactive species is proportional to the rate at which it undergoes
radioactive decay, stating the rate of radioactive decay rate for
an isotope in a given volume of water is equivalent to stating its
concentration. The units of radioactivity are disintegrations perunit time:disintegrations per second (dps), disintegrations per
minute (dpm) , and so on.
Since radioisotopes are generally measured by intercept ing the rad iat ions emit ted during the d is integration events, it proves more convenient to use activity rather
than mass units.

237

Incoming
Cosmic-Ray~
ProtonNeutronso Produced

'"

O'\-

Shatters.-:ornNucleus ofAtmosphericEntersNucleus ofAtmospheric
14N Atom,~

Net Result:.' @+.

•\ .+0N-14 + n(7p + 7n) + n

14 N Atom
Becomes
14C AtomoKnocks Out
a ProtonAmong the Fragments
Is a Neutron

\.

(a)

14 C

(b)

+P

(6p + 8n) + p(c)

Net Result:
14 C Atom Is Radioactive

Mean Lifetime:'T
M = 8200 ypars

Eventually
the t 4 C Atom.-·--'0

Spontaneously
EjectsHence:

n - e- -. p

an Electron•
Half Lifetime:'T

1/2= 5700 years~•(6 P + An) - e- ~ (7 P + 7 n)
14 C

(d)

(e)

Atom Becomes 14 N Atom(f)

Figure 5-1;, The "life cycle" of a carbon-14 atom.
Created in theatmosphere by the collision of a neutron (produced by pri~ary cosmic-ray protons) with a nitrogen atom, the average
4C atom"lives" for 8200 years.
Its life is terminated by the ejection ofan electron which returns the atom to its original form, 14 N•

238

leave the room until he has flipped ten heads in a row.
One student might be lucky enough to accomplish this feat in the firstfew minutes; another might not have succeeded even after several
hours of steady flipping.
From the laws of p~obability, howeve~,
we would predict exactly what the average success time would be.
The fraction of students present who would leave in a given timeinterval would be the reciprocal of this average success time.
If
there were enough students in the room when the flipping began, we
would observe that the number remaining would decrease in acco~dwith the same halving patte~n that holds for the radioactive decayof 14 C.
Although we could not predict the time required for any
specific indi vidual to make a successful run of ten heads, we
could predict the number of departures for a known time intervalrather closely.
Since there is not enough time for complete mixing during the
average lifetime of the 14 C atom, the more "remote" parts of thesystem will have less 14 C per unit of ordinary carbon than thoseparts "closer" to the birthplace of the 14 C atoms.
Since 14 C iscreated in the atmosphere, the atmosphere has the highest ratio of
14C to ordinary carbon.
Since the deep North Pacific Ocean is the
mos t remote part of the ocean, it has the leas t 14C per uni t ofordinary carbon. The 14 C/ C ratios that interest us most are those
for average surface and average deep water.
The difference in
14 C/ C ratios
between the mean carbon in these two water types
allows us to establish the rate of mixing across the main thermocline of the ocean.
If we want to make our estimates rigorous, we need a modelthat corresponds to the ocean in all its complexity. However, thefirst principles can be demonstrated by the simple two-box model
we used in chapter 1.
If we can understand how a hypotheticalocean consisting of two major water types - warm surface water andcold deep water - operates, we will have a basis for understanding
the methods used to treat the vastly more complic1ted problem
involving the real ocean.
This model as applied to
4 C is shownin f i gu r e 5- 2 •For the deep reservoir of our two-box ocean, three substances
must be conserved:
water, ordinary carbon, and 14 C.
We assumethat the surface and deep reservoir sizes and their carbon
contents have achieved constant values.
These reservoir sizes andthe distribution of ordinary carbon between them in no way betrays
the rate of mixing.
To find the rate we need a time clock.
Because 14 C is disappearing by radioactive decay, its distribution
between the surface and deep reservoirs does depend on the rate of
mixing!
-Thus we will wri te three equations: one for water, one forordinary carbon, and one for 14 C. The water conservation equation
very simply states that the amount of water going up, v up , must
equal the amount of water coming down, vdown.
We set both vup
and vdown equal to the parameter vmix:
V up= vdown = vmix

As stated in chapter 1, vrnix is thechanged between the two reservoirs.

239

5-1
rate at which water isex-WATER

RADIOCARBONCARBON

Surface

Surface

~

....
....
....

:.:

'.,

!

Vmix C su rface

V"'lixB

Vdeep C deep ~)
/
C
/

Deep

Deep

B + VmixCsurface= VmixCdeep

B = Vmix (C deep - Csurface)

A

deep

~

Deep

C*)

C*)

+B _

vmixCsurface _Csurfaceca) += vmix C deep C

deep

C

surfaceC*)

Vdeep C deep C

A

deep

5-2. Two-box model for the cycle of water, carbon, andradiocarbon (14 C) between the surface and deep sea.
If the system
is at steady state, then the amount of each of these substances
entering the deep reservoir must exactly match the amount lost.
Although the fluxes of carbon and 14 C from rivers and to sediments
should be included, they are negligible compared to the otherfluxes.
The straight arrows represent the fluxes of substances
carried by
wa ter;the wavy
arrow s,subs tances
carried byparticles. The dashed arrow represents loss by radioactive decay.

Figure240

Carbon is added to the deep reservoir in two ways:
it comes
down with the descending surface water and it falls from the surface in the form of particles destined for destruction in the deepsea.
The sum of these two contributions must be exactly balancedby the amount of carbon carried away by upwelling water'.
We canwri te the balance between these fluxes in the form of the equation: .
VmixCdeep = vmixCsurface + B

5-2

where Cdeep is the carbon content of upwelling deep water,Csurface is the carbon content of descending surface water, andB is the carbon added to the deep reservoir each year by the destruction of particles falling from the surface. Solving for B, we
obtain:B= Vrnix(Cdeep - Csurface)

5-3

A similar equation can be written for 14 C •Again, we have
two water fluxes.
First, 14C is carried up from the deep to thesurface wa ter and down from the surface to the deep wa ter as theresult of physical mixing.
Second, 14 C enters the deep water withdescending particles.
But there is still another factor to consider here:
one part in 8200 of all the radiocarbon in the deep
ocean undergoes radioactive decay each year.
We therefore need aterm that will give us the radioactive decay rate.
Rad iocarbon is measured in sea wa ter samples by extrac tingcarbon from the wa ter and measuring tha t carbon's radioac tivi ty.
This yields the amount of 14 C contained in the d issol ved carbon,
or the 14 C/ C ratio. * Labor'atory measurements thus do not givethe absolute amount of radiocarbon in the same water sample, but
the ratio of radiocarbon atoms \designated C*) to ordinary carbon
atoms (C).
So the &'Tlount of 1 C in descending sUr'face water isequal to the volume of descending wa ter vmix times the carbon
content of this descending water Csurface times the ratio of 14 C
a toms to ordinary C a toms in this carbon:g* ) surface
The amount of 14 C in upwelling deep wa ter is equal to vmix times
the concentration of carbon in deep water Cdeep times the ratioof 14 C atoms to ordinary C atoms in deep water:~* ) deep
Particles carry 14C down in almost the same proportion to othercarbon tha t is present in surface wa ter.
So the particle flux
14
*In the ocean-atrnos phere sys tern, for every
C atom there areabout 10

12

ordinary carbon atoms.

241will be the amount of carbon descending each year in particles B
times the ratio of 14C atoms to ordinary C atoms in surface water:

B vf"\* )C
surface
To calculate the radioactive decay of 14C, we first find the total
numbe r of 14 C atoms in. deep wa te r (by taking the volume of deep
water, Vdeeo' times the carbon content of deep water, Cdeep'times the ratio of 14C atoms to ordinary C atoms in deep watercarbon) and multiply it by the fraction of 14C atoms decaying eachyear, which we designate as A. Thus:
C* )AVdeep CdeepC
We can now write thereservoir as:
C
-

*)

C

deep

conservation equation for+

14C in the

deep

BC*
- )

surfaceC

surface5-4g* ) deep + AVdeep Cdeep g* ) deep
The conservation equation for carbon demanded that theparticle flux of carbon B be gi ven by \)mixCdeep\)mix
Csurface.
Thus we can eliminate B in equation 5-4 by substi tution. After combining terms, the result is:

Note that each term in this equation contains the concentration ofcarbon in deep water Cdee.
Therefore, this common factor canbe cancelled.
This is logfcal, since the rate of vertical mixing
cannot be dependent on the amount of carbon dissolved in the sea!
If we then sol ve the 14 C balance for the unknown mixing ra te\)mix' we obtain:
C* )

C
\)mix

AVdeep

deep

C

5-6

~* )

c* )
surfacedeep

We can express the volume of the deep ocean, Vdeep' in242

terms of the product of its mean depth, h, and its surface area,
Aocean. Equation 5-6 then becomes:
CC
AhAocean

*)

deep

C*)
C

surface

~* )

5-7
deep

Dividing numerator and denominator by C*/C)deep, we obtain:

5-8

AhAocean
Vmix

c*/c
C*/C

) surface-1) deepIn this e~uation, A is equal to 1/8200 years and h is 3000 meters(3.0 x 10
centimeters). Measurements of the radiocarbon content
of carbon extracted from sea water reveal that, as expected, there
is a deficiency of 14 C relative to C in deep water with respect tosurface wateri the 14C has not completely mixed with the ordinary
carbon. The 4C/C ratio in average deep water is about 12 percent
lower than it is in surface water. When we place these values inthe equation 5-8 and solve, we find that the yearly volume ofwater exchanged between the surface and the deep ocean is equal involume to a layer 300 centimeters thick with an area equal to thatof the ocean:
1

Vmix
Aocean= 300 cm/yr

1.12 - 1

Since the mean thickness of the deep reservoir is about 3000
meters, the transfer of a layer 3 meters thick between the deepand surface ocean each year yields a mean residence time of waterin the deep sea of 3000/3, or 1000 years.
A uni t of water (orsal t) has one chance in 1000 of returning to the warm surface
reservoir each year.

IMPLICATION TO THE DISTRIBUTION OF Ra-226
In the last chapter we discussed the distribution of 226 Ra in
the sea.
The statement was made that the ratio of 226 Ra to Ba
would not differ enough between surface and deep water to be useful as a crosscheck on the ventilation rate obtained from theratio of 14C to C.
Using the same model as [or l~., we can write
down equations for the balance of Ba and 22 Ra.
Since the major
addition of 226 Ra is to the deep reservoir, we shall write thesebalances for the surface reservoir.
For steady state for the barium in the surface reservoir:

243

5-9

Csurfacevrnix + B = Cdeep vrnix

where
Csurface
and
CdeeQ
are,
respectively,thebariumcontents of mean surface ana mean deep water, B is the particle
flux of barium from surface to deep water, and vmix is the watertransfer rate between the reservoirs.
For steady state for 226 Ra in the surface reservoir:
226 ) s Csurface vmix + 2 26 ) sB +
Ra/Ba
Ra/Ba
226
)
_226
)
A
Ra/Ba s Csurface VsurfaceRa/Ba dCdeep vmix+ P

5-10

where 226 Ra/ Ba )s and 226 Ra/ Ba )d are, respectively, the concentration ratios of these species in average surface and deep water,
A is the decay constant of 226Ra , Vsurface is the volume of thesurface reservoir and P is the input of 226 Ra from the sediments
and rivers to the surface reservoir. By simultaneous solution ofequations 5-9 and 5-10, B can be eliminated.
Also, P may be represented as follows:
226
P = f(A
Ra/Ba)s CsurfaceVsurface +226/
A
Ra Ba ) d CdeepVdeep)
where f is the fraction of the new 226 Ra
surface reservoir.
In this way we obtain:
226 Ra / Ba ) s
2 26 Ra / Ba ) dreaching

Vmix + fAVdeep

5-11
throughthe

5-12

Vmix + (I-f) AVsurface (CSurface )Cdeep

We can now replace vmix wi th the ratio of Vdeep/ Tmix, where
Tmix is the mean flushing time of the deep sea and A with
1/T226 Ra , where T226 Ra is the mean lifetime of 226 Ra atoms andobtain:5-13
2 26 Ra / Ba ) s
2 26 Ra / Ba ) d=

1 + (I-f)

Tmix
T226 Ra

Vsurface
Vdeep

Csurface
Cdeep

The maximum difference between the 226 Ra to Ba ratio of surface
and deep water is obtained if f = 0 (i.e, all the 226 Ra is .added
to the deep reservoir).
In this case:
226 Ra / Ba ) s
226 Ra / Ba ) dTaking1

1 + Tmix
T226 Ra

Vsurface/ Vdeep

5-14

Vsurface

Csurface

Vdeep

Cdeep

to

be

244

about

0.1,

Csurface/Cdeeptobe 0.3 and the value of Tmix to be 1000 years, we obtain:
226 Ra / Ba ) s

1

226 Ra / Ba ) d

1000

1 +

-2340

0.99
x 0.lxO.3

Thus, looked at in this way, the 226 Ra ratio difference between
mean surface and mean deep water should lie within the experimental error in measurements of 226 Ra/ Ba ratios (i.e., 3-5%).
It is not possible ~ priori to say what is the fraction of
the 226 Ra entering the ocean which does so via the surface.
Rivers contribute an amount of 226 Ra equal to several percent ofthat decaying in the sea. Shelf sediments constitute about 10% of
the sea floor.
Hence, were the flux of 226 Ra into the sea fromsediments everywhere the same, the proportion of 226 Ra entering
the surface sea would roughly match its proportion of the volume.
If we take as an upper limit that 15% of the 226 Ra entering thesea does so in the surface layer, then from equation 5-13:
1000
226 Ra /Ba) s

1 + 0.15
1 + (1-0.15)

mo
1000

= 1.050.lxO.3

2340
Again this ismeasurements.

close to the limit oferror on the

226 Ra/ Ba

ratio

DISTRIBUTION OF RADIOCARBON IN THE OCEAN
Before moving on to the other rates needed to understand theflow of chemical constituents through the sea it is important topoint out that in averaging all the 14 C data for deep ocean samples for use in the two-box model we have thrown out an enormous
amount of valuable information with regard to rates of mixing
within the sea.
Thus, we will pause here to take a look at the
distribution of 14 C in the world ocean.
Figure 5-3 shows vertical profiles of 14C/C ra tio* in the ocean as they looked prior*These ra t ios are exp ressed as the pe r m.i 1 d if fe rence from the
14 C/ C ratlo. .ln the atmosphere prlor.
to t h e onset 0 f
the'1n d ust~ial revolution (i.e., about 1850) and normalized to a constant
1 c/ 1 2c ratio. The equation used is as follows:lit~14C = o14 C - 2(o13 C + 25)(1 + ~ )

5-15

1000

14 C/ C ) sample [14 C/C)standard

14C/C)standard

245

J

1000

f:. 14 C (%f:.14 C (%0)-250 -200 -150 -100

o

-SO

o

0 )

-250 -200 -150 -100 -SOo....- ....- - . -....~-.....- .. O

-E 2
--.:

NORTH

:J: 3
I-

a..

27°N-54°W

-

2 E
--.:

ANTARCTICOCEAN

ATLANTIC
OCEAN

58°S-66°W

W 4

o

5

5
6 ' - -......-~-~-_ ....._ ..
f:.14C (%0)
o-250 -200 -150 -100 -50

o' - _. . . ._....._.....11-0_....._

E 2

--.:

-250 -200 -150 -100

J: 3
I-

a..

-50

0

o

/

NORTH

EOUATORIAL
I NOlAN

6~14C (%0)

--

-

...

OCEAN

PACIFIC
OCEAN

0°-Sl o E

35°N-152°E

-

2 E
--.:

W 4

o

5
6

5
,-_..a-_~_-..

__...._'""-----_......_ ....._ ......_ ..... 6

Figure 5-3~ 14 C/ C (expressed as Li 14 C) versus water depth at stations in the world's three major oceans and in the Antarctic.
The
progressive decrease in the values for deep water from Atlantic to
Antarctic to Indian to Pacific can be seen.
The arrow denotesthat portion of the upper water column contaminated with radiocarbon produced by nuclear tests at the times these samples were collected. This depth was determined from the tritium data for thesestations.
The dashed lines are the prenuclear Li 1 4 C trends forthis depth zone. These results were obtained on samples collected
as part of the GEOSEC.S program.
The analyses were made by Minze
Stuiver of the University of Washington and by Gate Ostlund of the
University of Miami (21, 22, 23).

246

to the influence of the testing of thermo-nuclear bombs (i. e. ,prior to 1957).
We make this point because a substantial amountof radiocarbon was produced during these tests.
As we shall discuss in chapter e this "anthropogenic" radiocarbon has penetrated
into the upper oceans.
The profiles in figure 5-3 show the trends expected fromthe distribution of oxygen and of the nutrient constituents.
Surface water carbon contains more l4 C than deep water carbon but
less than that in carbon from the preindustrial atmosphere.
Carbon in deep Atlantic waters has more l4 C than carbon from the deep
Pacific.
Deep waters from Antarctic and Indian Oceans have intermediate values.
The trend in radiocarbon through the deep ocean
is well displayed along the 4000-meter depth horizon (see figure5-4) .
As can be seen, the l4 C/ C ratios are highest at the northern end of the Atlantic and lowest in the northern Pacific.
Finally, the details of the natural radiocarbon distribution are
shown in a series of the vertical sections drawn from north tosouth through the major ocean basins (see figures 5-5 to 5-9).
These sections show that neither the new deep water formed at thenorthern end of the Atlantic nor that formed in the Weddell Sea
have the same l4 C/C ratio as seen in the warm surface ocean (seefigure 5-5).
The waters forming in the North Atlantic had (prior
to nuclear tes ting) a ~ l4 C value of abou t -70 0 /00 (compared toabout -50 0 /00 for warm surface water); those forming in the Weddell Sea had a value of -150 0 /00, which is only slightly higher
than that for ambient circumpolar water.
These differences reflect the fact that the source of deep water is mainly intermediate depth water which rises to the surface in the deep water
source regions.
These waters ar-e exposed to the atmospher-e only
long enough to be cooled bu t not long enough to exchange theircarbon with the atmosphere.
Another feature of interest is the penetration of the waters
flowing around Antarctica into the Atlantic above and below theNorth Atlantic Deep Water mass. This is well shown by the tonguesof low l4 C/ C water (see fipure 5-5).
The distribution of
4C is hence qui te consistent with what
we have learned of the distributions of the biologically cycledconstituents.
The oldest waters in the ocean have the least dissolved oxygen, the most nitrate, phosphate and silicate and thelowest l4C/C ratio.
In chapter 7 we will have more to say about
the time information carried by the distribution of l4 C/ C ratios
in the deep sea.

RATE OF CONTINENTAL RUNOFF
In chapter 1 we stated that the ratio of the rate of upwelledwater to runoff water entering the warm surface ocean was 30.
Now
that we have established the upwelling rate, let us turn our attention to the runoff rate.
In this cas e, no radioisotope isneeded. A very direct means is used:
the amount of water leaving
the mouths of the world's rivers is monitored and totalled.
We
can estimate the approximate magnitude of the result expected fromthis effort from two pieces of informa tion:
the average annual
rate of rainfall on the continents (-70 cm/yr) and the average

247

60

60

20

20

60

100

140

180

140

100 60

40

20

o

40

60 L....6.....0 ----2:uO--...I2'-O--6....
0~...;.L..1O....0--14~O~~~.-.,;~~~~~

LONG I TUDE

Figure5-4. Map show ing the dis tribu tion of 14C/C ra tios
(expressed as ~ 4C) at a depth of 4000 meters in the main basinsof the world ocean.
Based on resul ts obtained by Stuiver and by
Ostlund as -part of the GEOSECS program (21,22,23).

248

-

E 2

-:::r:
~

t-

o..
W
0

WESTERN
ATLANTIC

oooooooooo~14C

0000 0 000

4---.....&.-----......~---- .......------40
60
20
o
N
LATI TUDE

5L.---JL----L_ _........_-I.....a-.......

60
S

40

20

Ft~re 5-5. Vertical section of natural 14 C/C ratios (expressed as/), C) along the western basin of the Atlantic Ocean.
The values
shown are for natural radiocarbon (i.e., the distribution prior to
the influx of 14 C produced by weapons testing).
The thick dashedline designates the depth to which significant amounts of bomb 14C
has penetrated into the Atlantic at the time of the GEOSECS survey
in 1973 (as indicated by the depth at which tritium reached about
5% its surface water value).
The /),14 C contours above this line
are dashed because they are based on a ra ther limi ted number ofanalyses for samples collected prior to 1957.
Also shown by0
dotted lines are the °0 = 27.5 /00 (upper) and °4 = 45.2%0(lower) isopycnal horizons.
The results were obtained by Stuiver
and by Ostlund (21).
The tracks along which this and the sections
which follow were drawn can be found on the foldout map.

249

o........--...--~--y---y-- .......- .......--y----.
~

WESTERN
INDI AN\\........ - .. __

~

'1

1

:I:

_160

.....
a.
w

c_

~-180

/50_ 160
-110
~o
-180

W
q,

s~o

-170

-10

-I \

--i20'-

190

o~oo

q,

-190

°°0
%<1

o

o

20
LATITUDE

40
N

5 60
S

40

20
LATITUDEo

20
N

Figure 5-6. Vertical section of natural 14C/ C ratios (expressed in
14
11
C units) in the eastern Atlantic Ocean.
As in figure 5-5 thedark dashed line marks the depth to which significant amounts of
14C produced by nuclear testing had penetrated at the time of the
GEOSECS survey in early 1973.
Also shown by the dotted line is
the 04 = 45.2 % 0 isopycnal horizon.
The resul ts were obtainedby Stuiver and by Ostlund (21).

Figure 5-7. Vertical section of natural 14 C/ C ratios in thewestern Indian Ocean.
The contours terminate at the depth towhich significant amounts of 14C produced during nuclear testing
had penetrated at the time of the GEOSECS survey in 1978.
As wehave almost no data on samples collected prior to 1951 no attempt
has been made to reconstruct the contours of natural 14C/C ratios
in the upper part of this ocean. Also shown by the dotted line is
the 04 = 45.2 0 /00 is opycnal horizon.
The resul ts were obtainedby Ostlund
(23).

250

WESTERNPACIFIC[).14C
0/00

51--&_-L_..L.._L--I._--L...J...J.~_..L-l-~~-.;M.--:--:----J

60
S

40

20

0

LATITUDE

N

Figure 5-8. Vertical section of natural 14 C/ C ratios (expressed in
14b
C units) for the western Pacific Ocean from the Aleutian Arc at
the north to the Antarctic continent on the south.
As in theother 14 C ~ections the dark dashed line marks the depth to whichsignificant 14 C produced by nuclear testing had penetrated as of1974 when the GEOSECS Pacific survey was conducted. The limited
data 'tvailable for samples collected prior to 1957 suggest thatthe b 4 C values for surface water had a pattern similar to thatshown for the Atlant ic in figure 5-5.
Also shown by the dottedline is the a4 = 45.2 0 /00 isopycnal horizon.
The results wereobtained by Stuiver and by Ostlund (22).

251o........- ......-..--.-,,.-.---,.--r--,--..,-,.........

E 2

~

I

I(L

W
C

20

0
LATITUDE

20
N

Figure 5-9. Vertical section of natural 14 C/ C ratios (expressed in
14
11
C units) for a north-south section drawn through the eastern
Pacific. Once again the dark dashed line shows the depth to whichsignificant amounts of
14 C
produced by nuclear testing had
penetrated at the time of the GEOSECS survey in 1974.
Also shownby the dotted line is the 04 = 45.2 % 0 isopycnal horizon.
The
results were obtained by Ostlund and by Stuiver (22).

252percent of this water which survi ves evaporation and reaches thesea (35 percent).
With this knowledge, we can calculate that in
one year a layer of water 25 centimeters thick wi th the area ofthe continents will reach the sea. Since the continents occupy 30
percent of the area of the globe, if this water were spread overthe sea, it would make a layer about 10 centimeters thick (s1E$
figure 5-10).
This is very close to the result of 32 x 10
liters/yr obtained from the river gauges.
Since upwelling brings an amount of water equal to an oceanwide layer 300 centimeters thick and since rivers bring an oceanwide layer 10 cm thick, the ra tio of "'mix to "'runoff mus t be
30:
"'mix

300cm/yr Aocean"'runoff

10cm/yr Aocean

30

SEDIMENT ACCUMULATION RATESLet us turn our attention, then, to the ways in which sedimentation rates are determined.
All sedimentation rates in theocean are based on four methods of absolute dating.
One involves
l4 C, the same isotope we used to establish the rate at which theocean is mixing.
A second method invol ves the uranium daughter
products thorium-230 and protactinium-231, two radioactive isotopes produced by the decay of the uranium dissolved within thesea.
The third method involves the decay of lOBe an isotope whichlike 14C is produced by cosmic ray interactions in the atmosphere.
The fourth method involves radioactive potassium-40, which decays
to a gas, argon-40.
These four radioisotopic methods have all
been used to date deep sea sediments and, as we will see, theygive self-consistent results.
Deep sea cores contain unique stratigraphic markers which canbe traced over large regions of the sea floor.
Once a core containing such a marker has been dated, the level in all other deepsea cores where that stratigraphic marker is found can be assigned
this age.
Once the ages of several such horizons have been established, the determination of the distribution of accumulation
rates for the various components of marine sediments is greatly
simplified. There are three kinds of marks made in deep sea
s ediments that can be used for this purpose.
One is magnetic.
Reversals in the polari ty of the earth's magnetic field are recorded in' deep sea sediments.
The times when these transitions
occurred have been accurately determined by the potassium-argon
dating of volcanic rocks whose magnetism also records the polarityof the earth's field.
If we can establish which magnetic transition is which in a given core, ages can be assigned to theselevels.
The irregular sequence in time of these magnetic transitions (see figure 5-11) makes this assignment possible.
Another series of stratigraphic markers is based on faunalchanges.
For every group of organisms there are, over periods ofmillions of years, a few extinctions of old species, a few appearances of new species, and some unique changes in relative
abundances.
These changes also constitute key stratigraphic

253

I
Net Vapor Transfer+-+I

Net(P - E)
23 cm/yr

ATMOSPHERE

I
Precipitation(P)70 cm/yr..

Net(E - P)
10 cm/yr

I
I
I
I
I
I

1

t
Precipi tation (P)90 cmlyrI
Evaporation(f) I. . ..

.. .

47 cm/yr

1

I

I
I
I

I
I
I

I

I Evaporation(f)I

I

Runoff

100 cm/yr•+- 30 Percent of Earth Area (A edflh )1-+--------~~70 Percent of Earth Area (A earth)--'SEA
llTHOS

Runoff= Net Vapor Transfer

PH

E RE

I
I

I

...

.. . - .....

23 cm/yr X.3D Aearth cm 2;=:lO cm/yr X.70 Aedrth:--:',

Figure 5-10. The global wa ter cycle.

Abou t 10 percent more wa ter
evaporates. from the ocean than returns as rainfall.
The excess
falls on the continen ts as rain and snow and runs from thecontinents into the oceans through the world's rivers.

254horizons that, when dated, complement the use of magnetic reversals in assigning ages to various levels in a particular deep sea
core.
The final stratigraphic marker involves climatic changes.
The technique is less reliable because clima tic changes are repetitive at intervals of roughly the same duration.
In the main,
this method is helpful for the last two climate cycles.
Furtherback in time it is too easy to mismatch two curves and end up withthe wro ng age.
Howe ve r, as the las t magne tic reve r s aloe cu r red
700,000 years ago and as faunal methods are useful primarily froma few hundred thousand to a few million years ago, the clima tic
record is important for correIa tions over the las t 200, 000 years.
Climatic changes are recorded in deep sea cores in many ways, including the CaC0 3 content changes mentioned above.
18 0/ 160 ratiochanges in the CaC03 from the shells of benthic foraminifera have
proven to be the most reliable.
The reason is that these changes
are produced by the removal from (or addi tion to) the ocean ofwater isotopically depleted in 18 0 during the formation (or q~­
struction) of continental ice sheets.
The changes in 18 01 0
ra tio for sea wa ter so produced are spread uniformly through thesea.
Shells formed in deep wa ter prove superior because the deepsea is less sUb~ect to changes in temperature and river input
which produce 1 0/ 16 0 variations in shells formed in surface
water.
The 18 0 record from benthic forams taken from a core collected on the Solomon plateau in the western equatorial Pacific isshown in figure 5-12.

RADIOCARBON DATING
Now let us examine each of the dating methods in turn, beginning with the one based on radiocarbon.
We have already discussed
the way in which 14 C is generated in the a tmosphere and mixed wi th
the carbon in the ocean.
Foram shells and coccoli th tes ts whenformed in surface ocean wa ter contain 14 C a toms as well as ordinary C atoms.
Once formed, the shells constitute an isolated system; their 14 C content decreases with time in the logarithmic manner described previously:
half of the 14C disappears during thenext 5700 years, and so on.
The fundamental assumption of radiocarbon dating is that the
14 C/ C ratio at any place in the sea has always been the same as it
is today.
If this is so, then a shell sample from a deep sea core
yielding a 14 C/C ratio half that in present-day sea water carbon
must have formed 57qO years ago.
If we find the level in the core
where the ratio of 4 C/ C is only 25 percent of that in the present
surface ocean, the sample must have formed 11,400 years ago, andso on.
At an age of about 40,000 years, about 1 percent of theoriginal 14 C remains.
Beyond this point, the exchange of carbon
in the shells wi th their surroundings is thought to produce significant contamination and the method becomes questionable.
Al though the hal ving method is an easy way to envision wha t
is happening, it is only in rare cases that our sample will haveaged some mul tiple of 5700 years.
If we want to make ac tual age
calculations, we need an equation which relates the 14C/C ratio intime, or:

255

MATUYAMA~GAUSS
REVERSED EPOCHNORMAL

BRUNHESNORMAL EPOCH

JARAMILLO EVENT!'il'j!"! i'j!:il!I! I~li!i! ! ! ! !I ! ! ! jl[]~Io

1.........'.........1-,,-I--L.1_I.....L-.....L....-I-....I.-....

.5

OLDUVAI EVENT

lIB I

R

I

I

1

EPOCH

R

I

1.0

I I

I

I

1.5

I

I I

I

l...L..J.

2.5

2.0

MILLIONS OF YEARS BEFORE PRESENT

Figure 5-11. The time sequence of magnetic reversals as determinedby the potassium-argon dating of volcanic rocks.
The shaded areas
marked N (normal) indicate times when the magnetic field had thesame polari ty it does today, and the open areas marked R (reve rsed) indica te times when it had the oppos i te polari ty.
Longintervals' of time domina ted by ei ther normal or reversed polari ty
are called magnetic epochs and are named after scientists who madeimportant contribu tions to the field of magnetism.
Brief intervals of opposi te polari ty wi thin these epochs are called eventsand are named after the localities at which they were discovered.
V28-238+3.0 ++BENTHONIC
FORAMINIFERA

,,
,,
,

\

0-

\11,000 YRS~ +3.5

o

CD

ro

+4.0

\l

+

\ i/J\'*~ 1
\ t
\(
\/

+4.5 050"" 70,000 VRS"'126,OOOVRS "

+
100

150

250

DEPTH IN CORE (em)Figure 5-12. 18 0 /16 0 record from core V28-238 collected in thewestern equatorial Pacific.
Just over one full climatic cycle isshown. The large e 18 0 values are for glacial time and reflect thestorage of water deficient in 18 0 in the continental glaciers
which existed at that time.
Note that the cycle consists of a
gradual decline toward glacial conditions followed by a sharp
return to interglacial condi tions.
This record covers the las t
130, 000 years.
C.ycles of this type extend back over the last
million years.
These results were obtained by Nick Shackleton
of Cambridge, England (602).
The definition of the e 18 0 notation
is given on page 681.

256

)

x e- At
formation

5-16

The equation states that the 14C/ C ratio in the shell at any given
time t after formation is equal to the original 14C/C ratio timese (the base of the na tural logari thmic sys tern) to the power - At,
where A is the decay constant of 14 C (1 part disappears every 8200
years). Solving this equation for t, we obtain:( 14 C/C).
t
B200 In
formatlon
5-17(14C/C)fossil CaC03
The age t of the sample is equal to B200 years (the average lifetime of a 14C atom) times the natural logarithm (In) of the ratioof the 14C/ C ratio at the time of formation to the 14 C/ C ratio
found in the sample today.
We assume that ( 14 C/ C )formation is
identical to the ratio we find in present surface ocean water (andhence in living shells and coccoliths).
In figure 5-13 are shown radiocarbon ages made at series ofdepths in two deep sea cores.
Both show a surprising feature.
The age at the core top is not zero!
Ra ther the upper 7 centimeters of material in each core show a nearly uniform radiocarbon
age. Below 7 centimeters in each case the age begins to rise in a
regular way wi th depth as would be expected.
The cause of thesecore top anomalies is "bioturbation Tl •Organisms living on the sea
floor must "stir" the upper sediment much as worms living on thecontinents "stir" the soil.
Only when material has been buried
beyond a depth of 7 or so centimeters does it escape the stirring
action of these organisms.
The radiocarbon age of this bioturbated laye r is gi ven appro xima t ely by the ratio of its thi cknes s
to the sedimentation rate.
While not having any appreciable influence on ages obtained beneath the mixed layer or on sedimentation rates obtained from age versus depth plots beneath the mixed
layer, this process does have significance to some of the process es going on in the sea.
We have already deal t wi th one exampleof the role of bioturbation (i.e., its importance to the recyclingof short lived radioisotopes).
We shall deal with others in thechapters to come.

URANIUM SERIES DATINGThe s~cond dating method involves the element uranium and two
radioactive "daughter" isotopes produced by the decay of uranium,
230 Th and 231 Pa .
We will concentrate first on 230 Th , the productof long-lived 238U.
Uranium-238 was produced in the interiors ofmassive stars that sUbsequently exploded and dispersed theirproducts wi thin our galaxy.
It was incorporated into the solar
system when it formed from a galactic dust and gas cloud about 4.5
billion years ago.
Uranium-238 is unstable (that is, radioactive)
and has been gradually decaying since its stellar birth.
Since auranium atom has a me~n life of about 7 billion years, not enoughtime has elapsed for 238U to disappear and much of it remains to-257

14C AGE (I03 yrs )14 C AGE (1 0 3 yrs )
5

10

15

5
O
O-......-......

20

10

15

20

--....-----.,~-..,BIOTURBATEDLAYER (-7em)

-

-E

E

u

u

......"......"

:r:

::r:

I-

I-

w
o

w
o

a..

0.. 20LAST GLACI AL

LAST GLACIAL

3.5 em/I0 3 yrs

CORE ERDC92
PACIFIC OCEAN
2°S-157°E 1.6 km

CORE V19- 188
I NOlAN OCEAN
7°N-GloE 3.35km

70 '--_.......1._ _-.1-_ _....&-•

2.2 em/I0 3 yrs_

35 L . - _ - - L_ _.......

_

Figure 5-13. Radiocarbon ages as a function of depth in two deepsea cores.
The near constancy of these ages down to a depth ofabout 7 cm bears wi tness to the action of organisms which "burrow lT
into the sediment.
Below this zone of bioturba tion the ages show
the expected regular increase.
The changes in the slope of theage versus depth line at the time close to the end of the lastglacial period (i.e., ~10,OOO years ago) reflects the influence ofclima te on the ra te at which the various rna terials making up the
sediment rain onto the sea floor.
These results were obtained by
Pen6 of the Lamont-Doherty Geological Observatory (197, 199).

258

~~~

in our solar system.
As we discussed in the last chapter,U decays through a whole series of radioactive forms to stable
lead (206Pb).
Many billions of years from now all the 238 U will
have been converted into 206 Pb •Until then, we will have in nature a whole series of uranium decay products, produced as uranium
atoms and decaying towards lead (see figure 4-1).
The transition from 238U to 206 Pb requires the emission of 8alpha particles, each carrying away 2 neutrons and 2 protons.
Italso requires an equivalent number of electron emissions.
Thus,
a~ atom must go through many stages in the transition from 238 U . to26 pb , and the most interesting step, at least as far as datlng
goes, is its pause at 230 Th •This isotope has a half-life of
75,000 years (ideal for dating in the range of a ~~~ hundred thousand years).
Fortunately for dating purposes,
Th has a muchdifferent chemical cycle in the ocean than uranium does.
Uranium
has a mean residence time in the ocean of about 500, 000 years.
Its concentration is thus relatively high with respect to its abundance in natural rock.
On the other hand, as we have seen inchapter 4, thorium is extremely reactive in sea water; it is removed to the sea floor in a matter of a few decades.
This separation between parent 238U and daughter 230Th is the basis foruranium datin .
Recently formed deep sea sediments have a great
excess of 23 lfrTh atoms (over those supported by the decay of theparent 23 8U) .
The age of the sediments can be computed from thedisappearance after burial of these excess 230Th atoms.
However,
successful dating by the 230 Th method is possible only under special circumstances:both the rate at which the sediment particles
fall onto a given area of the sea floor and the rate at which the230Th atoms are incorporated into this falling sediment must remain constant with time. The trick is to develop criteria to showwhether these assumptions are valid for any gi ven core.
If thecore meets these cri teria, we use the 23 °Th content to calcula teages; if it does not, the measurements are of Ii t tIe value fordating.
Most of the sediments deposited in the open ocean do notaccumulate at a sufficiently constant rate to make reliable datingby this method possible. Fortunately some do.
In figure 5-14 is shown a plot of 23 °Th as a function of
depth in an ideal core that has a constant sedimentation rate of 2
cm/10 3 yrs.
If we measured the 230Th content of this material, we
would find that at a depth of 150 centimeters (or 75,000 years inage) the amount of excess 230 Th had dropped by a factor of 2 fromthat found at the top of the core.
If we then took a sample at
300 centimeters (150, 000 years in age), we would find only onefourth of the original excess 23 °Th content.
If we went down to450 centimeters, the age of the sample would be 225,000 years(three half-lives would have passed) and only one-eighth of theini tial should remain.
In such a core the 23 °Th concentration
should follow an exponential curve wi th depth.
In order to test
whether points fall along such a curve, it is best to transformthe coordinates of the graph in such a way that they lie along a
straight line.
This can be accom~lished by plotting the logarithmof the concentration of excess 20 Th atoms against depth.
If thecore meets the criteria of the method then the points should fallalong a straight line, as they do in figure 5-14.

259

64

>-

I-

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tU

<t
.s=
0

t-

",

C\I

en
en

c

G)

E
"'C
Cl>

C/)

0'

W

.......

W

0"'C

u
x

SED 1 MENTAT ION RATE
= 150 cm/75,000 yrs= 2.0 cm/I0 3 yrs

E

0::

«

W

Z
-I

1.5

3.0

4.5

3.0

4.5

6.0

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DEPTH(m)

)-

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64

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-

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et- c:
",

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(!)

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DEPTH(m)Pigure 5-14. The distribution with depth of excess 230 Th in an
ideal core with a constant sedimentation rate of 2.0 cm/l0 3 yrs
and a cons tant ra te of 23 °Th influx.
In such a core, the amoun t
of excess 230Th shows a drop by a factor of 2 for each interval of
150 centimeters penetrated (i. e. t the thickness of sediment isdeposi ted in one half-life of 23 Th ). The upper graph employ s alinear scale for the 230 Th content; the lower graph employs alogarithmic scale.

260

Figure 5-15. Actual 230 Th (a) and 231 pa (b) measurements for core
V12-l22 raised by the Columbia University research vessel Vema in
the Caribbean Sea.
The failure of the points to fall along asingle straight line in either graph reflects a combination ofexperimental uncertainty in the measurements (note error bars) andof nonideality in the core (the sedimentation rate and/or theisotope influx rate did not remain exactly constant with time).
The best fit solid line (a rate of 2.35 cm/IO 3 yrs) and dashedlines representing sedimentation rates 25 percent greater (3.00
cm/l0 3 yrs) and 25 percent smaller (1.90 cm/l0 3 yrs) are shown forboth isotopes.
Clearly with the observed scatter, the rate (andhence the age at any given depth) cannot be determined to better
than flO percent.
The agreement between the best fit rates for
230Th and 23 Ipa adds support to the validi ty of the ages derived
in this way.
The resul ts were obtained by KU, Goddard and Middelof the Lamont-Doherty Geological Observatory (611).

261

Figure 5-15 shows an actual set of 230 Th measurements on acore from the Caribbean Sea.
This core proves to be close toideal.
Although there is some scatter from a single straight
line, most of the points lie within twice the measurement error ofthis line.
The best fit line corresponds to a sedimentation rateof 2.35 cm/10 3 yrs.
To show the sensitivity of the sedimentation
rate to the choice of the best fit line, dashed lines representing
sedimentation rates of 3.00 cm/10 3 yrs (25 percent higher) and1.90 cm/10 3 yrs (25 percent lower) have been included for comparison. Even in the best cores, this method allows the average sedimentation rate to be determined with an accuracy no better than 10
percent.
The other product of uranium decay of interest for dating,
2.31 pa , is generated b~ the decay of a second isotope of uranium,
235 U.
Like 238 U, 23 U has a lon~ half-life (0.7 billion years).
Uranium-235 decays to stable 20 Pb, and like 238 U it must go
through many stages to make this transi tion to stabili ty.
Onestopping point is 231 pa , an isotope wi th a half-life of 32,500
years.
Like 230 Th , 231 pa is very insoluble and is removed fromsea water to the sediments within a century after it is produced.
We can measure the concentration of the isotope 231pa at various
depths in a deep sea core and plot these data on a logari thmic
graph in the same manner we did with 230Th.
If the points yield astraight line on such a plot, then the slope of this line can beused to calculate the sedimentation rate for the particular core.
Shown in figure 5-15 are plots of the 23 Ipa measurements for the
Caribbean core just discussed.
Since 231pa is less abundant incore tops and since it disappears more rapidly than 230 Th , it canbe measured wi th sufficient accuracy only to a depth of 3 meters(compared to a depth of 9 meters for 230 Th ). Like 230 Th , the best
fit line for 231 pa has a slope that yields a sedimentation rate of
2 . 35 cm/ 1 0 3 Yrs !
So, in this particular core, both is otopes tellus that the sediment has accumulated at roughly a constant rate.
Thorium-230 allows the mean rate to be established over an interval of 400,000 years; 231pa, over an interval of 150,000 years.
These isotopes have also been used to determine the growthrate of manganese nodules. Because these objects grow so slowly a
special sampling technique had to be devised.
Teh-Lung Ku did
this job by measuring off an area on the face of a nodule andscraping material from this area (224). By weighing these scrapings Ku could estimate the depth interval to which each successive
sample corresponded.
An example of the results
obtained in thisway is shown in figures 5-16.
As can be seen all three samples
taken from this nodule come from the outer millimeter. The slopesof the lines ioining the measurements gi ve growth rates of 4.0
mm/10 6 years ( 30 Th ) and 4.3 mm/10 6 ears (231 pa ).
Ku also measured the ratio of 34U to 238U activity in thesethree samples and in four others from deeper below the sur~~ce(see figure 5-17). As mentioned in chapter 4 the activity of
4U
in sea water is about 15% higher than that of 238 U•
The reason
for this excess is tha t during their creation in deep sea sediments 234 U atoms are knocked loose (by recoil frrnn the departing aparticle) from their positions in the clay minerals which carry

r.

262

>- >-

10 3 -

E

E

CL
"U

CL
"U

r10 2 >

r-

CL
"U....>>

~
E

~

~

2

r- 10
~

cf

y

",

N
(/)

en
w

U

x

U 0.1

«

0

0...
231 pa6104.3mm/10 yrs

I

2

::>

If')

N

N

en
w

(J)

en

U

lJJ
U

w

w

x

\

~

If')

(/)

w

1.0
0

-t->

lu
<l

1.0

DEPTH IN NODULE (mm)

x

.01

°

Figure 5-16. Plots of the logarithm of excess 230 Th and 231 pa
versus depth in a manganese nodule from the North Pacific Ocean.
The growth rates derived from these measurements are also shown.
These measurements were obtained by Ku while at Lamont-Doherty
Geological Observatory (224).

Figure 5-17. Plot of the logarithm of excess 234 U activity in thesame nodule for which 231 pa and 230Th diagrams are given in fi ure
9 Th
5-16. The slope expec ted for the growth ra te obtained using 20
is shown for reference.

263

the bulk of the uranium in these sediments.
These loosened 234 U
atoms have a better chance of migrating into the sea than do theirmore strongly bound 238 U brothers.
Nodules incorporate theiruranium from sea water, hence newly formed nodule material has 15%
more 234U than 238U activity.
Ku monitored the decrease of thisexcess wi th depth in the nodule and found that like 23 °Th and23l pa this decrease could be apRroximated as exponential.
Using
the half-life of 250, 000 for
4U the slope of the excess 23 4 U
t r end y i e 1d s an aceurnu I a t ion rat e consis ten t wit h t ha t for 2 3 aTh
and 231 Pa •fJIany nodules have now been studied in this way.
They allshow growth rates on the order of 2 to 5 mm/10 6 years.

BERYLLIUM DATING
Beryllium-10 is produced when the nuclei of atmospheric gases
are smashed apart during collisions with high energy cosmic
protons.
As this isotope has a half-life of 1.5 x 10 6 years it
has potential applications in the dating of ocean sediments andmanganese nodules.
The idea is that much of the lOBe produced in
the atmosphere is deli vered to the sea by rainfall and aerosol
impacts.
Like other metals these Be atoms find their way to ~hesediments and manganese nodules on the sea floor.
To harness I Be
for dating the same two requirements must be met as for the 230 Th
method; i.e., the accumulation rate of both lOBe and of the hostmaterial must remain constant with time.
TI1e application of the lOBe method was rather limited in thatits acti vi ty is very small.
Larger samples had to be dissol ved
and the beryllium isolated.
The lOBe in this beryllium then had
to be measured by low-level beta counting.
C:ounting times ofseveral days are required for each sample.
An e~ample of results
obtained in this way is given in figure 5-18.
As can be seen thethree points when placed on a logarithm activity versus depth plot
form a strai~ht line.
The slope of this line gives a growth rateof 2.4 mm/10 yr.
This and other lOBe results on nodules are very
important because they confirm the 23 °Th and 23l pa based growth
rates.
Because the latter were based on the outer millimeter ofmaterial many scientists were hesitant toaccept them.
As thelOBe results shown in figure 5-18 cover a depth interval of 1.5
centimeters they are not so easily disposed of.
Recently a new means of measuring lOBe has been found.
Instead of. measuring the beta decays of this isotope the atoms
themselves are measured in a high-energy mass spectrometer.
This
method requires only small samples (milligrams) and the measurements can be done quickly (a few minutes).
An example of resultsby this method is also shown in figure 5-18.
Again the growthrate obtained from lOBe is in agreement wi th that obtained fromthe uranium series measurements.
As summarized in table 5-1 thisis true for all the nodules studied to date by the lOBe method.

POTASSIUM-ARGON DATING
based

The time scale foron ages obtained

deep sea cores beyond 400, 000 years isby the potassium-argon method.
Potas-

264

100

100
2.4 mm/106 yrs........

E

Q.

-m

10

"0

/

0'

~

........

E

0.

10

"0
CI)

Q)

0

-

/

0'

~

4.5 mm/10 6 yrs

CD0

20 0 N
174°E

I

0

10

20

28°5
41°W

I

0

10

20

DEPTH IN NODULE (mm)

DEPTH IN NODULE(mm)Figure 5-18. The left hand graph shows the logarithm of the 1 DBeacti vi ty as a function of depth in a manganese nodule from theNorth Pacific Ocean.
These measurements were made by Ku and hiscolleagues at the Uni versi ty of Southern California using theconventional B-counting technique (260J.
The right hand graphshows the logari thm of the number of 1 Be atoms as a function of
depth in a manganese nodule from the South Atlantic Ocean.
Thesemeasurements were made by Turekian and his colleagues at Yale
University (262).

265

Table 5-=-t- Summa.ry of manganese nodule grcwth rates
by the 1 Ee method.
lOBe Flux

Location
104

atoms/em 2yr

GrCMth Rate
rmn/l0 6yr

as determined

Investigators &Institution

Ref.

400 N 1700 E

3.8

1.3

Shat et ale
PRL (India)(254)

21°N 11~W

5.0

6.3

Shat et ale
PRL (India)(254)

19°N 16c>W

1.4

3.0

Somayajulu(251)

Tata (India)

21oN 17~W

2.1

2.3

Sharma and Sanayajulu
PRL (India)(263)

21°N 17S'E

2.4

3.3

Sharma and Somayajulu
PRL (India)(263)

13°S 149°W

1.7

2.8

9°N 151°W

1.4-2

15°N 125°W

2.1

-2

Krishnaswami et ale
Yale University(264)

21°N 175°E

3.7

2.4

Ku et ale(260)

200 N 139°W

3.5

1.3

Ku et ale

(260)

2.7

4.5

Turekian et ale
Yale University( 261)

l28

00

S 41 W

266

Guichard et ale Centre(256)
des Faibles Radioactivities
Gif-Sur-Yvette (France)
Krishnaswami et ale
Yale University

u.s.c.

u.s.c.

(264)lsium-40 is another of the isotopes generated long ago in stars andlater incorporated in our solar system.
Its half-life is justover a billi&n years; so although it is also destined for extinction, some 4 K still remains in our earth.
Roughly 90 out of 100
40 K atoms
change themselves into 40 Ca ; the remaining 10 change
themselves into 40 Ar (see figure 5-19).
The proportion of 4U K
undergoing decay by each of these routes is a constant.
When agedeterminations are made, it is the argon produced by the decay ofpotassium that is measured rather than the calcium.
The reason
for this is that when a volcanic rock forms, gases are almost
totally excluded from the minerals that crystallize out of the hot
liquids.
Hence volcanic rocks are initially free of 40 Ar •With
time, the 40 K in these rocks undergoes radioactive decay and produces argon which remains trapped in the crystal lattices of therock.
If we measure the amount of 4 Ar and the amount of 4 OK in
one of these rocks, it is possible to calculate the length of timethat has elapsed since the rock crystallized.
To apply this method directly to deep sea sediments requires
the chance presence of a layer of volcanic ash.
The normal Kbearing minerals found in sediments are detrital in origin andbear the great age of the rocks from which they were eroded.
Unfortunately, volcanic ashes in deep sea sediments are quiterare, so the potassium-argon method does not prove to be very
profitable for the direct dating of sediments.
The most powerful application of the potassium-argon method
was in the establishment of the chronology for magnetic reversals.
Volcanic rocks believed to have formed in the last few million
years were collected from allover the earth for this purpose.
For each sample, the polari ty of the magnetic field at the timethat the rock crystallized was established by measuring the magnetic vector in an oriented sample of the rock.
This is possible
because the iron-bearing minerals in basaltic rocks were magnatized in accord with the direction of the earth's magnetic field at
the time the rock cooled.
If a rock cooled during an interval
when the magnetic field was reversed, then its magnetic vector
will point toward the south pole instead of the north pole.
Roughly half the rocks collected fell in the north pole category
and half fell in the south pole category. The potassium-argon ageof each rock was then carefully measured.
The results showed that
all the rocks with ages of less than 0.7 million years were normal
in polarity and that most of the rocks dating between 0.7 and 2.5
million years were reversed polari ty, except for those formed inseveral time intervals during which the earth's magnetic fieldbriefly returned to normal (see figure 5-11).
The same temporal sequence of magnetic polarity observed in
volcanic rocks exists in deep sea cores.
The orientation of iron
bearing grains settling to the sea floor is in accord with theearth's field.
Thus the ages
obtained from the volcanic rockrecord can be applied to the reversals seen in deep sea cores.
Ofcourse, for a core to be dated in this way, it must penetrate atleast the youngest of these ma.gnetic reversals (i.e., that at
700,000 years). Magnetic correlation dating allows dating back toabout 4,000,000 years.

°

267

40K

(19p + 21 n)

90 PERCENTELECTRON EJECTION(nO~ p+ + e-)

10 PERCENTELECTRON CAPTURE(p++ e- ~ nO)

40

40

Ar

Ca
(20p + 20n)

(18p + 22n)t 1/2 = 1.1 X 109 Yrs

Figure 5-19. The

scheme for radioactive decay of the long-lived
One nucleus in ten transforms bypotassium isotope,
40 K.
capturing an elec tron, the other nine transform by emi tting anelectron.
The argon isotope, 40 Ar , produced by the electron
capture process provides a means of dating volcanic rocks.

268

AGREEMENT AMONG DATING METHODS
We will cite two examples of the agreement obtained when several dating methods are applied to a single core. The first core,
RCII-209, was taken by the Columbia University research vessel
ROBERT CONRAD in the equatorial ocean just north of the equator.
One sample from a depth of 11 centimeters yielded a 14 C/ C ratio 25
percent the ratio in surface water, indicating that the sample had
aged for two half-lives of 14 C (that is, for 11,400 years).
Thi~measurement gives a sedimentation rate very close to 1. a cm/IO
yrs.
In this core, the first magnetic boundary occurs at 700
centimeters.
As this reversal has been dated by the potassiumargon method to be 740,000 years in age, the average sedimentation
ra te was again 1.0 cm/10 3 years.
By measuring the varia tion in
CaC03 content with depth in the core, it was possible to establish
the climatic record and, particularly, the point at which the second sharp warming before present occurred.
In other cores, theage of this boundary has been determined by the 230 Th method to be
125,000 years.
The use of this time marker yielded a sedimentation rate for the core of about 1.2 cm/10 3 yrs.
So the threemethods - direct 14 C dating, the magnetic correlation of an event
dated by the potassium-ar~on method, and the climatic correlationof an event da ted by the
30 Th method - all give roughly the same
resul t!
Tne second core, V12-122, is the Caribbean core already used
as an example of 230 Th and 231 Pa dating (see figure ;-15).
These
two methods yielded concordant ra tes of 2. 35 cm/IO
yrs.
At 9
meters depth in this core, there is a faunal change tha t can becorreIa ted from core to core.
This paleontologic event has beenda ted at 400, 000 years by interpola ting between the core top andthe level of the first magnetic reversal (740 ,000 years) in 5
cores for which a magnetic stratigraphy has been established.
From this age we can obtain a sedimentation rate of 2.25 cm/l0 3
yrs.
In the same core, 14 C ages of 8950 years at 27 centimetersand 18,300 years at 57 centimeters were obtained.
These ages
suggest a sedimentation rate of 3.0 cm/l0 3 ~rs.
Thus, a~ain rates
of sed imenta tion based on 4 ° A r, 23 °Th,
31 pa ,
andC dating
methods are reasonably consistent.

COMPARISON OF MODEL AND OBSERVED RATES OF caC03 SOLUTION
We can use the two- box oc ean mod el to pred ic t the ra te a twhich CaCO~ dissolves in the deep sea.
As we discussed in chapter
1, it is possible to show, from the variations in alkalinity andtotal dissolved carbon from water mass to water mass, that forevery 4 moles of C in the form of organic molecules oxid ized in
the deep sea roughly 1 mole of C is dissolved in the form of
CaCO 3.
The total amount of carbon returned to solution per unit
area in the deep sea is given by:
B = vmix(Cdeep - Csurface)

5-18

Hence the amount of CaC03 dissolving per unit area is given by:269

Table 5-2. calcite accumulation rates during interglacial t~es for cores taken above
the lysocline (the ~C:03 entry in the table shCMs extent of supersaturationof the water at the depth the core was taken).
Core
No.

Lat.

Long.

Water
Depth
kIn

280
1665
1681
V29-192K
V30-10lK
V22-226

35°N
26°N
26°N
5lfN
44°N
300 N

440Ttl
45°N
45°N
17°W
32°W
45°W

4.30
2.90
3.97
2.37
3.30
3.13

A240-MI

15°N
17°N

68°W
7lfW

4.18
2.80

0°

3.75
3.32
3.87

V12-122

Bulk
Frac.
CaC03*
Accurn. Rate caco 3 Accum. Rate2
Wl/kg
cnv'10 3 yr
% w'cm 10 3 yr
NORm AfJT ",AN'"rl"C
- 2
60
3.5
1.5
0.84
-25
0.6
85
-10
0.6
0.77
83
3.5
1.7
65
2.. 2
1.4
77
1.2
-25
1.1
95

IlC03

CARIBBEAN
-10
-28

Rei".

(211)
(+)
(+ )
(+ )
(+ )
(+)

2.3
2.4

56
64

0.9
1.2

(213)
(229)

ATLANTIC
3.1
2.3
2.0

68
(85)
88

1.6
1.8
1.6

(196)
(195)
(+)

1.3

84

LO

(+)

EQUA'IDRIAL PACIFIC
-15
1.5
3.6
- 5

86

zr

1.2
0.5

(199)
(+ )

~UA'IDRIAL

-22
-31
-20

Al8o-73
A18O-74
V30-41K

CO

0°

23°W
21f'W
23°N

V22-72

29°8

35°W

3.12

ERDC

RC20-16

2°S
6°N

157°E
86°w

1.60
1.97

Vl9-54
Vl9-56
Vl9-61

17°S
17°S
17°S

114°W
11SOW
116°W

2.96
3.15
3.42

SOU'lli PACIFIC
-12
1.25
1..3
- 5
0.44
- 7

60
78
90

0.5
0.9
0.4

(235)
(+)
(235)

V19-188
RCI4-35

7°N
OOS

61°E
90 0 E

3.36
3.02

EQUAIDRIAL INDIAN
2.2
- 5
2.6
- 5

82
85

1.6
2.0

(197)
(+)

RC9-150
RC17-69
V18-207

31°S

llSOE
33°E
87°E

2.70
3.31
2.45

70

2.1
1.1
La

(+)
(+ )

SOUIH ATLANTIC-40

SOU'IH INDIAN
3~S

26°s

-10
- 3
-12

*Mean for all cores in this table:+Unplblished L9mont data.

3.7
3.8
1.0

1.2 g/cm 210 3yr

270

48

96

(+ )

1

5-19

5
We have also shown that vmix is equal to a layer of ocean water
300 centimeters thick per year (or ~ 0.3 kg/cm 2 yr).
The concentration of dissol ved carbon in deep water averages 2.25 x 10- 3
moles/kg.
In warm surface water it averages 2.00 x 10- 3 moles/kg.
We find, then, that CaC0 3 must be falling from the surface
ocean into the deep ocean and dissolving at the rate of 1/5 x 0.3
kg/cm 2 yr x 0.25 x 10- 3 moles/kg, or 0.015 x 10- 3 moles of carbon
per square centimeter of ocean surface per year.
Accumulation
rates of material in deep sea sediments are not normally expressedin these units, but rather in units of grams of CaC03 per squarecentimeter of sea floor per 1000 years.
So we need to convert
f rom one set of uni t s to the othe r.
It happens that 1 mo Ie of
CaC03 weighs almost exactly 100 grams.
Thus 0.015 x 10- 3 moles/
cm 2 yr of CaC0 3 is equivalent to 1.5 g/cm 2 10 3 yrs of CaC03, or:g

Calcite accumulates on about 20% of the sea floor and dissol ves from the other 80%.
As summarized in table 5-2 the mean
accumulation rate of calcite for cores from above the lysocline is
1.2 g/cm 210 3 yrs.* If the rain rate of calcite onto the other 80%of the sea floor is also assumed to be 1.2 g/cm 2 10 3 yrs then the
total rate of solution of calcite is 0.80 x 3.6 x 10 18 cm 2 x 1.2
g/cm 210 3 yrs or 3.4 x 10 15 g/year.
The average calcite so-lution
rate per unit area of sea floor becomes 0.95 g/cm 210 3 yrs.
To this rate of calcite solution we must add the contributionof aragonite. From sediment trap studies it is estimated that therain rate of aragonite is about 20% that of calcite. As virtually
all the aragonite reaching the deep sea dissolves, the solution
rate of aragonite is 0.2 x 1.2 or 0.25 g/cm 2 10 3 years.
Addingtogether the solution rates of aragonite (0.25 g/cm 210 3 yrs) andcalcite (0.95 g/cm 2 10 3 yrs) we get a total "observed" solution
rate for CaC0 3 of 1.2 ~cm2103 yrs.
This compares favorably withthe value of 1.5 g/ cm 10 3 yrs obtained from the two-box model.
Hence, th~ rates of ocean mixing based on the 14 C within the ocean
is consistent wi th sedimentation rates based on 14C measurementson sediments!
SUMMARY

In this chapter we have considered the means by which theabsolute rates of oceanic processes are ~stablished. The rate of
2 3*This value can be compared with those of 0.9 g/cm 10
yr, 0.42
2
3g/cm 10 3 yrs and 0.4 g/cm /l0
yrs obtained by Honjo from his
sediment trap deployments at sites, respectively, in the equatorial At lant ic, the temperate North At lant ic and the North Pacific(see table 1-3).

271

river influx is estimated from river gauge records.
The rate ofvertical mixing is derived from the difference between the 14 C/ C
ratio in average surface and average deep water.
The rate of
CaC03 accumulation is obtained by absolute age determinations madeusing radioisotopes found in deep sea sediments.
A cross-check
was conducted by comparing the observed CaC03 solution rate withthe rate predicted from the two-box model of oceanic mixing.
The
agreement of these two estimates shows that the estimate of therate of oceanic mixing is
consistent with the rate at which
calcite is accumulating on the sea floor.
The fact that deep sea sediments have accumulated at similar
rates over the last few million years (based on the potassiumargon calibrated magnetic reversal chronolo~~), as ov~r the past
few hundred thousand years (based on the 2 Th and 2 Ipa dating
methods), and as over the past few tens of thousand of years(based on 14 C dating), suggests that the present mode of ocean
operation is not atypical.
The mill has been grinding at more orless the same rate for at least 1,000,000 years!

PROBLEMS

5-1

Planet X has an ocean with a stratification similar to ours
but wi th a mean depth of 8000 meters.
Measurements revealthat its surface water has a 14 C/ C ra tio twice that in its
deep water.
What is the average residence time of water
beneath the thermocline of this ocean?
Assume the mid-depthof the thermocline to be 800 meters.

5-2

What cosmic ray production rate (atoms/cm 2 min) was needed tobalance the decay of 14 C on the surface of the earth prior to
nuclear testing? The needed data are as follows:
Reservoir
Atmosphere

Amount of Carbon
10 16 moles

14 /C Catm

5

295

0.83

Living biosphere

5

0·97

Soils.

100.90

IVlarine sediment7

o. 9 5Ocean

The mass of CaC03 in marine sediments is taken to be equal tothe amount in superlysocline sediment down to one mean lifeof 1 4 C ~i • e., 0.2 x 3. 6 x 10 1 8 cm 2 x 8.2 x 10 3 Yrs x l . 2 g
CaC0 3 /cm 10 3yrs x 10- 2 mole/g).
The preindustrial specificactivity of atmospheric C02 was 13.5 dpm/g C.

5-3

Sediment taken at a depth of 50 centimeters in a deep sea
core contains foram shells with a 14 C/ C ratio 12 percent thatfor sediment from a depth of 10 centimeters.
What is the272

apparent sedimentation rate?

5-4

The following results were obtained using uranium series
measurements at various depths in a deep sea core. The units
are disintegrations per minute per gram of sediment (dpm/g).
Depth
em

238 U
dpm/g

234U
dpm/g

230 Th
dpm/g

2
22
42
62
82
102
122

1.3
1.5
1.4
1.2
1.5
1.2
1.4

1.4
1.6
1.3
1.3
1.5
1.3
1.5

65.5
33.6
17.4
9.6
5.7
3.5
2.4

What is the apparent sedimentation rate?
If the 23 1pa
activity is 6.9 at 2 centimeters, what should it be at 42
centimeters?

5-5

The first change from a normal to a reversed magnetic field
is found at 14 meters depth in a core.
If the sedimentation
rate has been roughly constant, over what depth range could
the 14C method be used to reliably date this core? Over what
depth range could the 23 0Th and 231 Pa methods be used toobtain average sedimentation rates?

SUPERPROBLEM#5After learning about Theft's four-reservoir aquariums twogeochemists
specializing inradiocarbon measurements,
MeaslyStover and Cyote Eastland, decide to use the same approach to
explain to their students why radiocarbon is distributed as it isin the ocean.
They quickly abandon the idea of using an actual
aquarium, however, because while Theft could scale all the dimensions and rates to match the oceanic proportions, the half-life ofradiocarbon cannot be scaled.
Thus they decide to recreate thefour-reservoir model in a computer.
Stover and Eastland ;ncrease
Theft's volumes (see figure 1-19) by a factor of 1.3 x 10 1 (i.e.,
the total, volume becomes "'1.3 x 10 21 liters) and Theft's flow
rates (see figure 1-19) by a factor of e x 10 14 (i.e. the transport of water from the surface reservoir above Atlantis to Atlantis becomes 4 x 10 15 liters/day or 1.5 x 10 18 liters/yr).
Theysimulate Theft's filters and backflushers by the appropriate computer manipulations.
They delete the input and output fluxes andassume that phosphorus is entirely recycled.
They fix the C/Pratio in the reservoir system as a whole to be 1000/1 and assume
that the particles carrying P to the deep reservoirs have C/Pratios of 100/1.
They then run the model until the phosphorus andcarbon distributions come to steady state.
If the average P04
concentration (i.e., that obtained were the waters from all four
reservoirs thoroughly mixed) is 2.4 x 10- 6 moles/liter then what273

are the P04 and Ee02 concentrations found by Stover and Eastland

for the individual reservoirs?
Stover and Eastland then adJust the alkalinity of their sea
water to yield a pea
of 280 x 10- 6 atmospheres 1n the surface
reservoirs.
As they ~o not allow CaC03 to form and include no N03
1n their model the alkallnl ty 1s uniform among the four reservoirs.
~e atmosphere
over the model 1s also given a Pe0 2 of
280 x 10- atmospheres.
They set \he interface areas of the two
surface reservoirs to be 2.4 x 10 1 square meters (for Indopac)
and 1.2 x 10 14 square meters (for Atlantis).
The C02 exchange

rate across these interfaces 1s set at 20 moles/myr.

The amountof CO 2 1n the atmosphere comes out to be one-sixtieth of that 1n
the model ocean.
To keep things simple Stover and Eastland assume that no
carbon isotope fractionation occurs rcross the air-sea interface
and during photosynthesis (thus the 4C/C ratio in falling particles is just the same as that in surface water,. They also assume
that no soils or forests exist (thus all the ~C atoms are either
1n the atmosphere or dissolved in the sea).
Stover and Eastland then add -1 ~C atoms to their model atmosphere at the rate of 2 x 10 2.6 atoms/yr.
The~ let the computer
run through a period of 40,000 years so that
4C decay comes tosteady state with 14 C production.
During this time the distribution of l~C among the 5 reservoirs comes to steady state.
Whatvalues do Stover and Eastland get for the 14 C/ C ratio in each of
the ocean reservoirs
relative tothat
1n the atmosphericreservoir?
How would these results differ if the ocean were
lifeless?

WOOFER THE TRACER DOG

274

ekapte/t 6
WHAT KEEPS THE SYSTEM IN WHACK?
CONTROL MECHANISMS OPERATING IN THE SEA

INTRODUCTION
In chapter 1 we stated that any given constituent must, on
the time scale of its residence in the sea, be removed from thesea at nearly the same rate as it is added to the sea. We did not,however, consider the means by which the ocean maintains this balance. This will be the major task of this chapter.
The general principles governing the constituent balances inthe sea are similar to those governing economics.
Natural mechanisms are at work which tend to drive the rate of output of any
constituent toward equality with the rate of input. We will referto these mechanisms as negative feedback controls.
The adjective
negative refers to the fact that they work to eliminate any imbalances.
The principle is simple.
Let's say for some reason therate of loss of a given constituent were to become greater than
its ra t e of input.
Th en the am 0 un t 0 f t his con s tit u ent s tor e din
the sea would begin to fall (i.e., its concentration throughout
the sea would become smaller).
Somehow this lower concentration
would lead to a reduction in the rate of output of this constituent from the sea.
Athough we will discuss the factors controlling the concentrations of the major anions and cations in the sea, our discussions in this chapter will emphasize the controls operating on the
biologically important constituents of sea salt (i.e, P, Si, C, N
and 02).
In these discussions we will use the two-box model ofchapter 1 and in so doing will circumvent the complications stemming from variations in the concentrations of these properties
among the deep ocean basins.
At the end of this chapter, we willconsider wh~t geological events might have pushed the system outof whack and were such events to have occurred how they might berecorded in the sediments.

PHOSPHATE CONTROLS
We start with phosphate for two reasons.
First, its control
mechanisms are more easily understood than, for example, those fornitrate or for total dissolved carbon.
Second, phosphate is theultimate limiting nutrient.
A reexamination of figure 1-5 reminds us that the following
processes are important to the budget of phosphate in the sea (atleast within the context of our two-box ocean).

275

1.

the rate at which phosphate is supplied to the sea by
rivers;2.
the rate of transfer of water between the surface anddeep sea;
3. the fraction of these residues returned to solution in
the deep sea.
If any of these three parameters were for some reason to change,
the balance between phosphorus input to the oceans and loss fromthe oceans would be dis rupted.
Feedback controls would become
operative and would restore balance.
Let us take a hypothetical example, where one of these parameters is changed and the others are kept constant. Specifically,
let us assume that the rate of transfer of water between the surface and deep sea were suddenly to double.
The situation before
and after this doubling is shown in figure 6-1. For simplicity wehave neglected the small flux of phosphate with downwelling waterand have assumed 100% utilization of this consti tuent in surface
water.
Prior to the change the system is assumed to have been inbalance, the flux of phosphorus to the sediments just balancingthat from ri vers.
As a first approximation we will assume thatafter the change all the fluxes wi thin the ocean would double.
Each year twice as much phosphate would reach the surface ocean,
twice as much phosphorus would sink in organic residues, twice asmuch phosphorus would be returned to solution in the deep sea andtwice as much phosphorus would become entrapped in sediments.
However, the amount of phosphate entering the sea would remain thesame.
Hence, the ocean would lose more phosphorus each year thanit gained and the phosphate content of the deep sea would begin to
decrease.
As the phosphorus content of the sea fell all the fl uxes
wi thin the sea would fall by corresponding amounts.
As shown in
the figure 6-1, when the phosphate concentration has reached 0.75
its ini tial value, the misma tch between input and loss will drop
to half the value which existed just after the mixing rate was
dOUbled.
When the phosphate concentration of deep sea water had
reached half of its ini tial value the mismatch would have disappeared altogether.
All the phosphorus fluxes would have returnedto their initial values.
How long would this readjustment take?
Just after the rateincrease two percent of the ocean's phosphate would be lost to
sediments each 1000 years.
This yields a response time of 50,000
years (and a half adjustment time of 50,000 x In2 or about 35,000
yea r s ) .
Ttius, th e ph 0 s phate con c e n t rat ion wou I d dec reaset 0 0. 75
its initial value 35,000 years after the mixing rate was doubled.
The change of deep ocean phosphate ion concentration with time isshown in figure 6-2.
The approach to balance would be asymptotic
and would go on for many half-times before becoming imperceptible.
Similar imbalances could be generated by halving the river
input rate or by doubling the fraction of organic residues preserved.
Table 6-1 summarized the critical information for theseother scenarios. In each case the imbalances in the phosphate budget are remedied by a negati vee feedback loop operating between
the concentration of phosphorus in the deep sea and the productivity of plant matter.
The adjustment time (in thousands of years)

276

I

~
L:.

1

1

316.8
(P04]=2.4XI0-

6

475.2

m~~es

C

BEFORECHANGE(P04]= 1.8XI0-6m~~es

AFTER

35,OOOyrs

3.2

4.8

1

I

r

633.5

I

316.8

6
(P04] =2.4 x 10- m~~es

0

AFTER
CHANGE[P04]=1.2 X10-

6

m~~es

BALANCEREESTAB.

6.4

3.2

Figure 6-1. Response of the phosphate cycle to a change in rate oftransfer of water between the deep and surface reservoirs.
Panel
A shows the steady state before the change occurred; output
matches input.
Panel B shows the situation immediately after thetransfer rate is doubled.
All the fluxes within the sea are close
to twice what they were before the change; loss exceeds gain.
Panel C shows the situation 35, 000 years after the change.
Thephosphate content of the sea has been reduced to three quarters of
its initial value.
The excess of loss over gain has been reduced
to half what it was just after the mixing rate was doubled.
Panel
D shows the situation several hundred thousands of years later
when balance has been restored.
All the fluxes are the same as
they were originally.
The phosphate concentration in deep wateris, however, only half as large.

277

.--..-

2.4

C"

~

-"0

INITIAL

2.1

STEADYSTATE

E::l 1.8......,a.

FINAL

STEADY
STATE

TIME
OF

lJJ
lJJ

c 1.5
..---.
v
0a. 1.2C I RCU LATIONCHANGE

t

'-'

-70

0

70

140

ELAPSED TIME (THOUSANDS OF YEARS)-

Figure 6-2. Phosphate concentration in the deepfunction of time after the change in mixing rate.

278

reservoirasaTable frl. Effect of changes in the operational parameters (i .e., deep sea ventilation rate, river input
rate and preservation fraction for organic particulates) for the two-box model on the fluxes and contentsof phosphate in the ocean. The times are in thousands of years.

QUANTITY
Mass in sea
Cone. in deepRiver input
Upw. Flux
Part. Flux
LDss to seds
QUANTrrrY

INIT.
VALUE

t

10 15 moles
10- 6 moles/kg
10 10 moles/yr
10 10 moles/yr
10 10 moles/yr
10 10 moles/yr

3.2
2.4
3.2
316.e
320.0
3.2

2.4
3·2
2.4
1.8
3.2
3.2
633.6 475.2
636.8 478.4
6.4
4.8

UNrrrsINrrr.
VALUEt

10 15 moles
10- 6 moles/kg
10 10 moles/yr
10 10 moles/yr
10 10 moles/yr
10 10 moles/yr

3.2
2.4
3.2
316.8
320.0
3.2

2.2
3.2
2.4
1.8
1.6
1.6
316.8 237.6
318.4 239.2
3.2
2.4

UNITS

INIT.
VALUE

t

3.2
2.4
3.2
316.8
320.0
3.2

2.4
3.2
2.4
1.8
3.2
3.2
313.6 235.2
316.8 238.4
6.4
4.8

UNIrrSN.........

ex>

t = 0

HALVE MIXINl RATE
t = 140 t = 2eO t = ex>

1.6
1.2
3.2
316.8
320.0
3.2

3.2
2.4
3.2
158.4
161.6
1.6

4.8
3.6
3.2
237.6
240.8
2.4

ill lliLE MIXING RATE

= 0 t = 35 t = 70 t =
2.0
1.5
3.2
396.0
399.2
4.0

HALVE INPur RATE=

0 t = 70 t = 140 t =

00

t

5.6
4.2
3.2
277.2
280.4
2.8

6.4
4.e
3.2
316.8
320.0
3.2

IDillLE INP or RATE
t = 140 t= 0 t = 70

=

00

~

Mass ill sea
Cone. in deepRiver input
t.pw. Flux
Part. Flux
Loss to seds
QUANTITY
Mass in sea
Cone. in deepRiver input
Upw. Flux
Part. Flux
Loss to seds

10 15 moles
10- 6 moles/kg. 10 10 moles/yr
10 10 moles/yr
10 10 moles/yr
10 10 moles/yr

=

2.0
1.5
1.6
198.0
199.6
2.0

1.6
1.2
1.6
158.4
160.0
1.6

4.8
3.6
6.4
475.2
481.6
4.8

00

t = 0

HALVE FRACTION PRES.
t = 140 t = 2eO t =

1.6
1.2
3.2
156.8
160.0
3.2

3·2
2.4
3.2
15e.4
161.6
1.6

4.8
3.6
3.2
237.6
240.8
2.4

roillIE FRACTION PRES.
0 t = 35 t = 70 t =

2.0
1.5
3.2
196.0
199.2
4.0

6.4
4.8
6.4
633.6
640.0
6.4

3.2
2.4
6.4
316.8
323.2
3.2

5.6
4.2
6.4
554.4
560.8
5.6

5.6
4.2
3.2
277.2
280.4
2.8

00

6.4
4.8
3.2
316.8
320.0
3.2

is given by the reciprocal of the fraction of the ocean phosphate
lost during a period of a thousand years.

SILICATE CONTROLSThe situation for silicate is quite similar to that for phosphate.
The main link between input and loss is through the productivity of plants.
Within the context of our model the rate ofsilicate loss to sediments is proportional to the rate at whichsilicate reaches the surface ocean.
Thus in the simplest of all
worlds, were the rate of ocean mixing to double, the amount ofsilica raining toward the sea floor would also initially double.
This would certainly increase the amount of opal buried each thousand years in the sediment.
This would lead to a deficit in thesilicate budget which would in turn cause the silicate content of
the deep sea and in turn the flux of opal toward the sediments tofall.
Eventually, the silicate budget would come back into balance.
There is one important difference between the cycles of silica and phosphorus.
The fraction of opal dissol ved would alsochange.
As we have shown in chapter 2, the preservation probabili ty of opal depends on the rain rate of opal to the sea floor.
The higher the rain rate, the greater the preservation probability.
This higher preservation probability would further enhance
the rate of opal
storage.
Thus the imbalance between silicate
supply and burial would be even larger than those caused by the
production rate change alone.

CARBON CONTROLS
We have seen that carbon is removed from the ocean both in
the form of organic residues and in the form of calcium carbonate
hardparts.
This fact complicates the system of controls for thisconstituent.
If we assume for the sake of simplicity that theorganic residues being stored in marine sediments carry carbon andphosphorus in their 105 to 1 biochemical proportion, then the situation can be rather easily sorted out.
Under these circumstances, a fixed fraction of the carbon will be tied to the phosphorus cycle and will leave in the form of organic residues.
Theremainder will then be available for loss as calci teo
The fraction tied to the phosphorus cycle is given by the ratio of C/F inorganic residues (i.e., ~l05) divided by the e/P ratio in river
input. The ~atio of carbon atoms to phosphorus atoms reaching thesea from rivers is about 500.
Thus, about 20% (100 out of 500) of
the incoming carbon atoms leave the sea as organic residues and
80% (400 out of 500) leave the sea as calcite.
Our problem is todetermine what keeps the accumulation rate of calcite in tune with
the input of "excess" carbon (i.e., that beyond the amount needed
to match phosphorus in organic residues).
The first point to be made here is that both calcium ions andcarbonate ions exist in surface water in amounts far beyond thatneeded by marine organisms to form their CaC03 cages.
As we sawin chapter 2, the surface ocean is highly supersaturated (i. e. ,about five fold) with respect to the mineral calcite.
Thus, there280

is no way to impose a feedback control on the amount of CaC03
formed per unit of phosphorus upwelled. Marine plants and animals
can form as much CaC0 3 as they need.
The control must rather involve the fraction of CaC03 preserved in marine sediments.
In today' s ocean marine organisms generate about four times
more calci te than is being preserved in marine sediments.
About
three quarters dissolves (the basis for this estimate is given intable 6-2). It is in this solution process that the control lies.
This is easily seen as follows.
If calci te storage in sediments
were to exceed the "excess" carbon made available then both thealkalini ty and 1:C02 content of the sea would begin to drop.
Thealkalini ty would drop at twice the rate (in ~eq/kg) as the 1:C02
concentration (in ~m/kg). This would cause the difference between
alkalini ty and 1:CO 2 and hence the CO 3= ion concentration to decrease. A decrease in C03= ion concentration in the deep sea
would expand the domain of undersaturated water and thereby reduce
the fraction of calci te preserved in the sediment (see figure 63). Eventually this would restore the carbon balance!*
We have already shown in chapter 1 that the time constant for
the replacement of carbon atoms in the ocean is about 180, 000
years. The time constant for readjustment of the CaC03 budget is,however, much shorter than this.
Rather than being equal to thetotal amount of carbon dissolved in the sea divided by the removal
rate of carbon to sediments, the response time of the C03= ionconcentration to changes in the budget of CaC0 3 is equal to thetotal amount of carbonate ion in the sea divided by the removal
rate of calci te to marine sediments.
Since only about 1/30th ofthe carbon in the sea is in the form of C03=' the adjustment
time for the CaC03 cycle is only 6000 years!

INTERACTIONS BETWEEN THE PHOSPHATE AND CARBON CONTROLSWith -these facts we have the wherewithal to consider simultaneously the impact of a change in operation of the ocean systemon both the balances of total dissolved carbon and phosphate. Let
us return to our original example (i.e., what would happen if therate of water transfer between the deep and surface reservoir of
our two-box ocean were to double?).
Ini tially this would throwboth the balances for C and P out of balance. Both organic residues and calcium carbonate would accumulate at too rapid a rate.
This would cause the P0 4 concentration in the sea to fall (andhence also the ra te of production of organic residues and theirassociated' CaC03).
It would also cause the CO 3= ion content of
the deep sea to fall (and hence also the fraction of calcite which
dissol ves) . Because the time constant for the adjustment of the
CaC03 budget (~6000 yrs) is more than an order of magnitude
smaller
than
that
for
the
adjustmentofthephosphorus*As there are 100 Ca++ ions in the sea for every CO 3= ion, on ashort time scale it will be the C03= ion concentration which shifts
rather than the Ca++ ion concentration.
For times longer than a mill ion years (i. e., longer than the res idence time of ca lc ium in thesea) both shifts in Ca++ and in C03= ion concentration would have to
be cons idered.

281

Table 6-2. Preservation of calcite for various depth ranges in theworld's
maj or oceans.
The las t column gi ves es tima tes of thefraction of the global calcite production preserved in theparticular zone.
The sum of the entries in this column give theglobal preservation factor for calcite.
Fractionof Calcite
Preserved

Atl.

Area
(10 14 m2 )Ind. Pac.

0.0 to 0.2
0.2 to 2.e
2.8 to Lysocl.
Transition Zone
Red Clay Zone

-0.10*-0.20*
-0.90
-0.10
-0.02

0.14
0.22
0.35
0.21
0.14

0.03
0.11
0.18
0.14
0.30

0.10
0.21
0.18
0.14
1.15

0.27
0.54
0.71
0.49
1.59

0.01
0.03
0.18
0.01
0.01

Total

-0.24

1.06

0.76

1.78

3.60

0.24

Depth Range
km

Contribution to
Preservation
Factor+
Total*These percentages are low because much of the sediment in these
depth ranges contains sizable amounts of organic matter.
The
consumption of this organic matter by benthic animals and bacteria
and worms releases C02 to the pore waters and hence dissolves
CaCO 3.
+Obtained as follows:
Area in Depth Rangex Fraction Calcite Preserved
Area of Sea Floor282

[C0 3](fLmol/kg)
100

50

150

•

••

\•...
•

~

E

\...
•
\:•

2

~

~

~

W
0

4

•,e

•

71\

J: :3

a..

•
•
••

LYSOCLINE ••••IF [C0 J••
DOWN BY 15%:••(2.8km)•3

TODAY·S
LYSOCLINE(3.55 km)\:•

LYSOCLI NE~IF[C03]
•

:\
•

•

•

5

UP BY 15 %(4.3 km)Figure 6-3. Plot of carbonate ion content (solid curve) and calcite saturation carbonate ion content (dashed curve) in the western South Atlantic.
Today' s lysocline is located at the depth
where these two curves cross. The dotted'curves show hypotheticalsituations· where the carbonate ion content of the deep sea israised and lowered by 15%.
As can be seen the resul ting changes
in lysocline depth are substantial.

283

budget (- 100,000 yrs) the mismatch between the rate of entry to
the sea of "excess" carbon and the calcite accumulation rate would
be remedied more rapidly than would the discrepancy in thephosphate budget.

NITRATE CONTROLS
We have already stated that the ni trate cycle is probably
tied to the phosphate cycle through ecologic feedbacks.
If theratio of the amount of nitrate in the ocean to that of the amountof phosphate were for some reason to fall below the ideal biochemical ratio of about 15 then nitrogen-fixing organisms living insurface water would be given an ecological advantage over thoseorganisms which cannot fix nitrogen. This would lead to a restoration of the biochemical ratio.
On the other hand, if the ocean
were to find itself with an excess of N over P then the non-nitrifying organisms would gain an advantage pushing the ratio back toward the ideal value.
The time constant for the adjustment of the NIP ratio in thesea is not known.
The reason is that neither the rate of fixationof N2 (to form N0 3 -) nor the rate of denitrification of N03 (toform N 2) is known.
We saw in chapter 3 that roughly 0.1% of the
N03 is lost to N2 0 during each mixing cycle (i.e., each thousand
years).
Were this the only mode of nitrate destruction, then thecurrent adjustment time of the NIP ratio would be about 1,000,000
years.
However, we have already seen that substantial amounts ofnitrate are being utilized for oxidation of organic matter in waters free of2.
'We have also seen that there is evidence for NO 3
utilization (i. e., reduction to N 2) in large regions of the deep
Pacific where O 2 is present in the water column.
In this situation the utiliza tion is thought to occur wi thin fecal pellets asthey fall through the water column or after they reach the sea
floor.
As no one has found a way to quantify the importance of
these processes it is currently not possible to give a time scale
for the adjustment of the nitrate balance in the sea.
Our assumption will be that ni trate balance can be achieved on a shortertime scale compared to that required for phosphate.
This remains
however, an unproven assumption.
As such it constitutes one of
the key missing pieces in the marine chemistry puzzle.

°

DISSOLVED OXYGEN CONTROLS
So far in this discussion we have assumed the fraction of organic residues destroyed in the deep sea to be a constant. Clearly, this fraction must vary with the condition of the system. One
obvious dependence would be on the oxygen content of deep sea water. Observations of the composition of sediments show that thosebathed in anaerobic water have far higher organic carbon contents
than sediments bathed in aerobic waters. The reason is apparently
that the organisms which use N03 and S04 as oxidizing agents andthe organisms capable of fermentation (i.e., breaking organic
material down into CH 4 and CO 2) are all single celled and hence
incapable of the complexi ty of biochemical manipulation employedby 02 utilizing organisms.
Thus, they less efficiently use the284

available food and leave more residues.
Were the oxygen content
of the sea to fall the fraction of the sea floor bathed in anaerobic waters would as a consequence expand, and so also would thefraction of organic residues preserved.
Thus, there is a tie between the cycles of phosphate and dissolved oxygen.
The ratio of the amount of 02 in the atmosphere (plus ocean)to the net rate at which O2 is being produced gives a replacement
time for atmospheric oxygen of about 10 million years.* Thus, forthe short times we have been discussing, the oxygen content of theatmosphere can be considered to be constant.
Do the changes in the cycles of carbon and phosphorus inthemselves lead to a change in the 02 content of deep ocean water?
The answer is yes, bu t the si tua tion is complicated.
We can see
what goes on by considering our favorite example. What would happen to the 02 content of the deep sea if the water transfer rate
between the surface and deep sea were doubled?
Initially there
would be no change.
Al though the amount of organic residues
reaching the deep sea would double, so also would the amount ofoxygen carried by the downwelling surface water.
One fundamentalof marine chemistry is that for any gi ven pattern of ocean operat.ion the amount of oxygen in deep sea water depends only on theratio of the P0 4 concentration in the deep sea to the partialpressure of oxygen in the atmosphere and not on the rate at which
the deep sea is ventilated.
The following hypothetical experiment will help the reader tograsp this concept.
A bucket full of deep sea water is placed in
the sunlight on a winter day and is allowed to remain there until
plants have converted all the P0 4 (and N03) to organic residues
and until the O2 content of the water has achieved equilibriumwith the atmosphere.
All the organic matter is then removed fromthe water (and the sides of the bucket) and is placed in a small
vial. The bucket is taken to a dark room, at which point the organic matter is returned to the water and the bucket is sealed sothat there is no air space above the water.
It remains this way
until bacteria have consumed all the organic matter.
If the water
initially had the average composition of deep sea water then it
contained about 2.25 llm/kg of P04.
Plant matter containing thisamount of phosphorus would have an 02 demand of about 300 lJ1ll/kg.
The amount of O2 dissolved in the water when it was taken insidewould have been about 350 llm/kg.
Thus during storage in the darkroom the oxygen content would drop to about 5011m/kg. This outside
to inside qycle could be repeated over and over again. At the endof each exposure the resul t would be the same.
As long as thebucket was stored outside long enough for the plants to utilize
the limiting nutrients and inside long enough for the bacteria to*As t h·1S net 1S. a very .sm~ll ~raction of. the rate at which 02 is produced by photosynthesIs 1t 1S not est1mated by subtraction of theworldwi~e :espir~tion rate from the worldwide photosynthesis rate.
Rather It IS est1mated from the rate at which organic carbon accumulates in marine and continental sediments.
The
small net productionof 02 by photosynthesis is largely matched by corresponding oxidationof carbon, iron and sulfur passing through the erosion-deposition
sequence.
Thus, it need not lead to a rise in atmospheric 02.

285

consume the organic matter then the result would be independent of
the rate at which the bucket was cycled.
Now let's say that after one of the exposures to the sun and
air the organic material sieved from the water were di vided into
two equal parts.
Half is put aside and half is put back into thewater in the dark room.
In this case thez content would drop
during storage in the dark room by only 150 ~m/kg (to 200 ~m/kg).
Were the cycle repeated under this new condition only half as muchorganic matter would be produced and consumed each time (because
half of the phosphorus and ni trogen would remain in the organic
matter put aside in the small vial).
The ratio of [P04J (after
exposure in the dark room) to [0 zJ (after exposure to the air)would be only half as great as before and hence the 02 demand (indarkness) would be only half as great.
Let us consider yet one more perturbation of this hypothetical system.
The half of the organic matter put aside is restored
to the water to bring us back where we started.
Then instead of
placing the bucket outside on a cold winter day we place it outside on a warm summer day.
Again, plants would grow and fully
utilize the nutrients and again the water would become saturated
with atmospheric Oz.
However, the Oz content of the water would
be only about 200 ~m/kg. This being the case, when the bucket was
taken to the dark room and sealed from the air the Oz would be
consumed before the organic debris in the water had been consumed.
This would then permi t organisms capable of using NO 3 as an oxidant to take over.
The wa ter in the bucket would have become
anaerobicl
These bucket experiments provide quite a good analogue to theocean.
The exposure to sun and air of course is equivalent to thetime spent by water in the surface mixed layer.
The separation
and readdition of the organic matter is equivalent to the settlingof organic fragments from the surface mixed layer to the deep sea.
The organi-c matter and water follow different routes to the deepsea where they are recombined.
The d.eep sea is dark and is isolated from the air as was the bucket in the dark room. We approximated a change in the P0 4 and N03 contents of the sea by removing
half of the organic matter from the bucket cycle. We approximateda change in the atmospheric oxygen content by changing the airtemperature over the bucket.
The point of all this is to demonstrate that the oxygen content of the bucket after storage in darkisolation was controlled by the ratio of phosphate to dissolved
oxygen and not by the frequency at which the bucket was moved backand forth between the two environments.
Now back to our consideration of perturbations in the ocean's
dissolved oxygen cycle.
Immediately after the deep to surface
mixing rate was changed, the Oz content of deep sea water would
remain unchanged.
However the amount of O2 in the deep sea would
then begin to rise as the P0 4 content of deep sea water fell.
Theless P0 4 in the sea the smaller the amount of organic rich residues formed per unit of water upwelled and the smaller the demandfor 02 in the deep sea.
After the sea had reestablished the balance between loss and gain of phosphorus it would have half theoriginal P0 4 content.
The average AOU (i.e., saturation minus insitu oxygen content) for the deep sea would be half its initial°286

value.
A smaller area of ocean floor would be bathed in 02 freewater.
Also 02 would invade more deeply into fecal pellets andinto marine sediments permitting scavengers to burrow more deeply
into the sediment. Thus the fraction of organic residues preserved would likely be smaller.
In this sense our treatment of thephosphorus cycle was oversimplified in that we failed to take into
account the change in the 02 content of deep water and hence inthe fraction of organic matter preserved.

MAJOR ANION CONTROLS
What controls the makeup of anions in the sea? We have seenthat chloride is the most abundant anion, that sulfate is the second most abundant, and that bicarbonate runs a very poor third.
As these three anions do not find happy homes in igneous rock,when liquids and gases from the interior reach the surface andcrystallize, the chlorine, sulfur, and carbon they contain arelargely. released ei ther in volatile form to the atmosphere or indissolved form to the ocean.
The place of residence on theearth's surface for these three elements is either in the sea itself or in chemical precipitates formed within the sea and buriedin the sediments.
Inventories of the amounts of these chemical
precipitates show that roughly one third of the chloride currently
resides in sodium chloride (NaCl) deposi ts in sediments and tha t
about three-quarters of the sulfur is present in sulfate and sulfide minerals in sediments. By contrast only about 0.1 percent of
all of the carbon released from the earth's mantle is now dissolved in the sea (as shown in table 6-3 the remaining carbon,
about 1000 times more, resides in sediments as limes tone, dolomite, and organic residues).
Have the proportions of these elements contained in the oceanchanged with time?
At present, the ocean is far from saturated
with respect to any salt capable of precipitating chloride (Cl-)
ion or significant quanti ties of sulfate (30,+ =) ion.
If all thesodium chloride and calcium sulfate in sediments were dumped backinto the ocean they would readily dissolve. Thus one might wonder
why all this salt has not dissolved in the sea.
The answer liesin the observation that salini ties considerably higher than theocean average are generated in isolated arms of the ocean, likethe Persian Gulf, where water is evaporated more rapidly than itis added to the basin by rain and by runoff from adjacent lands.
In some of these basins, the salinity rises to the point where
CaSO,+ and 'even NaCl precipi tate.
We suspect from the geologic
record that most of the very large salt deposi ts found buried in
sedimentary sequences were formed in arid peripheral basins which
did not communicate rapidly with the sea.
Once such an evaporite
deposit has formed, it can be preserved for hundreds of millionsof years if it is buried under a layer of impermeable shale.
Eventually, continental drift and the resultant crustal upheavals
bring these salt deposi ts back to the surface where fresh water
can dissolve them.
The contained salt is then dissolved and sent
back to the ocean.
From this discussion we can see that the fractions of these
two major anions present in the ocean at anyone time depend on

287

Table 6-3. Distribution of the volatile elements among the atmosphere, ocean, and sediment reservoirs(679a). Amounts are given inmoles.

Hydrogen

Element

CarbonChlorine

Ni trogen

SulfurATMOSPHERE
Form

H 2O

CO 2

-

N2

Amount1 x 10 18

5 x 10 16-

3 x 10Form

H2O

HC03+C0 3=

Amount8 x 10 22

3 X 10

Form

H2O

CaC0 3+MgCa(C0 3 )2
+Kerogen*Amount"'8 x 10 21

"'3

X

10

880 x 10 20

"',30

X

20

-

OCEAN

18

Cl-

N0 3-

S04=

7 x 10 20

3 x 10 16

4 x 10 1 9

NaCl

Kerogen

SEDIMENT

21

"'4

10 20

"'11

X

20

"'1

10 20

CaS04
FeS2 + FeS-1.5 X 10 20

10 20

4 X 10 20

2 X 10 20

10

X

TOTAL
AmountX*Kerogen refers to the residues of organic matter found in
sediments.

288

chance tectonic factors.
How often is a large basin formed thathas just the right geometry (and climate) to extract salt rapidly
from the ocean?
How often does a salt deposit thus formed comeback into contact with fresh waters and dissolve? At times in thepast when peripheral basins in arid areas extracted salt more rapidly than usual or when old deposits were being destroyed lessrapidly than usual, the salt content of the ocean must have decreased.
When the reverse was true, the salt content of the sea
must have increased.
The salinity of the ocean has therefore
probably fluctuated over the epochs of geologic time.
Such fluctuations, however, have probably been fairly small because notenough salt is stored in the sediments to effect very large salinity changes.
For example, if all the NaCl and CaS04 now known toexist in sediments were dis sol ved in the sea, the salini ty of theocean would only increase by about a factor of 1.5.
The fluctuations that have occurred must have been very gradual.
From the rates at which these processes are currently taking
place, we estimate that a CI- ion remains in the ocean for perhaps
200 million years before being "captured" in a basin where NaCl is
actively precipitating. The CI- ion may reside in the sediment soformed for another 100 million years before returning to theocean.
Thus a typical Cl atom has spent about one third of its
time in the sediments as NaCI and two thirds of its time in theocean as CI- ion. Sulfur, on the other hand, must spend a greater
fraction of its time entrapped in sediments.
A typical S atom remains in the ocean several tens of million years before it isfixed into gypsum (CaS04) or iron sulfide (FeS or FeS 2).
This S
atom then remains in the sediments for 100 or so million years before it is released by solution and carried back to the ocean.
Thus we see that the salinity of the ocean is determined by
two factors:
the ratio of the elements hydrogen*, chlorine, andsulfur in volatiles released from the earth's interior and thetectonic fa~tors influencing the partitioning of sulfur and chlorine between the ocean and the sediments. The earth received theseelements in a particular ratio from the solar nebula, and it is inroughly this ratio that they appeared at the surface. Since theirarrival, the earth's geography has dictated their distribution between solid NaCI, solid CaSO 4, solid FeS or FeS 2, and dissol ved
sea salt.
MAJOR CATION CONTROLSThe ocean must remain electrically neutral.
That means thatonce the anion content of the sea is fixed, the cation content is
automatically determined. For each negative charge due to CI- and
S04 = ions, a corresponding positive charge must exist.
A greatmany cations could conceivably contribute to this balance.
However, only four elements - sodium, potassium, calcium, and magnesium - are both sufficiently soluble in sea water and abundant incrustal material to make significant contributions.
Just as
S04 = and Cl- contribute the major part of the negative charge,*Virtually all the hydrogen on the earth's surface is now, andalways has been, in the ocean as water.

289

Table 6-4. Comparison between the chemical composition of sea
water and the chemical composition of average river water for the
major cations in sea salt.
Also given are the corresponding
residence times in the sea.

Element

Mean
Sea Watermoles/m 3

Mean
River Water
moles/m 3

Na
K
Mg
Ca

470
10
53
10

~0.10*

0.04
0.16
0.36

*Corrected for sea salt.

290

Ratio
Sea
River

Residence

-5000
-250
-330
-30

-200

T~me

10

yrs

~10

-13
~1

Na+, K+, Ca++, and Mg++ contribute the major part of the positive
charge.
Whereas the major anions in sea salt are found in only
trace quantities in igneous rocks, the major cations are present
in lar~e quanti ties.
Continental erosion therefore car~ies more
K+, Mg +, Na+ and Ca++ ions to the sea than it does S04- and Clions. How then does the sea manage to maintain the required electrical balance?
Somehow the ocean must dispose of the cation excess which it continually receives from rivers by precipitating
new minerals similar in chemical composition to those being weathered on the continents.
If the amount of cations contained by thesea were to become a little too high (compared to the amounts of
Cl- and S04 = ions), the OH- content of sea water would rise.
This in turn would lead to a more rapid formation of the mineralphases that remove these cations.
Some such self-balancing mechanism must exist to maintain equality between cations and anions
wi thin the sea.
Unfortunately, we are only beginning to learn
about the alumino-silicate mineral phases being formed in the sea.
Until we know more, arguments such as the above must remain vague.
What controls the proportions of these four cations in sea
water? Let us begin by examining their relative abundance in river water and, in turn, in the rocks undergoing erosion.
In riverwater (see table 6-4), we find that for every mole of K there areabout 2.5 moles of Na, 4 moles of Mg, and 9 moles of Ca.
In theocean, for every mole of K there are 50 of Na, 5 of Mg, and only 1of Ca.
Rivers therefore carry these ions to the ocean in muchdifferent proportions than are currently present in the sea.
The
reason the sea stores these cations in proportions, which differ
so radically from the proportions found in river water, is thatthei r probabili ties of removal from the sea are qui te different.
The sea rids itself of calcium most readily.
This element is aconstituent of the calcite produced in great quantities by marineorganisms. After Ca, potassium removal is the most rapid. Apparently (we ~re not yet certain) K is taken up on the detrital alumino-silicate minerals brought to the sea by wind and water.
By
comparison, the sea has considerable difficulty getting rid of thesodiwn and magnesium it receives. The only way it can do so is tolet such very high contents of Na and Mg build up relative to K
and Ca that the minerals which form and the volcanic rocks which
are altered on the sea floor are literally flooded with sodium andmagnesium.
The ratios of the four major cations in the sea thus depend
upon the ratios in which these cations are being added from rivers
and the difficulty the sea has in ridding itself of them.
The
former is fixed by the composition. of continental rocks. The latter is fixed by the affini ty of the minerals forming wi thin thesea for a particular cation.
For each unit of "difficulty" theocean has in getting rid of Ca, the sea currently has 10 units of"difficulty" in getting rid of K, 13 units of Hdifficulty" in getting rid of Mg, and 200 units of HdifficultyH in getting rid ofNa.
An analogy might help here.
Four teams are competing in agame of musical chairs.
The players walk around a circle until
the music stops and then everyone dives for a chair.
As there aremany more participants than there are chairs, not every player

291

gets a seat.
The first team is comprised of professional athletes, the second of high school students, the third of housewives, and the fourth of retired office workers.
Clearly, in thefirst round, the athletes and the students will get the greaternumber of sea ts.
Now, contrary to the normal rules in such agame, those who succeed in grabbing chairs step aside and let
those who failed try again.
In the second round, the housewives
and the retired office workers will fare considerably better
agains t the diminished ranks of their opponents.
Obviously, in
our analogy, Ca and K are the athletes and the students of theocean and Mg and Na are the housewives and retired office workersof the ocean.
The sites in the lattices of forming minerals are
the chairs. As the Ca and K ranks are depleted, the less reactive
Mg and Na atoms find it easier to capture their share of latticesites.
A comparison between the cation composition of the ocean andother saline water bodies in closed basins is revealing in thisconnection.
As shown in table 6-5 despi te major differences intotal salt content and anion composition, these saline lakes have
Na/K and Mg/Ca ratios ranging over only a factor of ten.
The
Na+K/Mg+Ca ratios, on the other hand, cover an enormous range,
sugges ting that the singly charged Na and K ions compete for adifferent set of mineral lattice sites than do the doubly charged
Mg and Ca ions.
In our analogy, it would be as if the retired office workers and the students were competing for chairs of one
color and the housewives and the athletes for chairs of anothercolor.
The rivers feeding the saline lakes have chemical compositions reasonably similar to the average composition of thoserivers feeding the sea.
Apparently any water body that must ultimately dispense the ions it receives to its sediments (rather than
through an outlet to the sea) achieves a Na/K ratio roughly 25
times and a Mg/Ca ratio roughly 25 times higher than the respective ratios in incoming river water.
These ratios suggest that incompeting for sites accommodating singly charged ions, Na is only
about one-twentyfifth as effective as K, and that in competing for
doubly charged ion sites, Mg is only one-twentyfifth as effective
as Ca.
There are three major routes by which the cations Na, Mg, andK leave the sea.
First, they are being precipitated within thesea or at the sediment water interface as silicate minerals.
Second, they are precipitating within the sediment as authigenic
minerals.
Third, they react with the basalts beneath this sediment cover.
Are there ways in which we can quantify these processes?
Recent studies of the hot waters is suing forth from hydrothermal vents dotted along the crest of the East Pacific Rise provide a breakthrough wi th regard to the study of the interaction
between sea water and the hot basalt being emplaced along ridge
axes.
As shown in figure 6-4 chemical analyses of waters issuingforth from these vents are not the same as those for bottom water.
Further, when the concentration of any given element whose concentration in hydrothermal water is anomalous with regard to that inbottom water is plotted against temperature of the water, a strong
correlation is found.
The hotter the water, the larger the292

Table 6-5. Comparisons of the cation compositon of sea water withthe cation composition of various inland seas. See reference 547
for data sources.
Water
Body
Dead Sea(Israel)Desmet(U.S.A. )Bolac(Australia)
Caspian(USSR)
Tagar( USSR)
Timboram(Australia)
Gnarport(Australia)
Ocean
Murdeduke(Australia)
Great Salt
(U.S.A.)
Corangamite(Australia)
Pyramid(U.S.A. )Walker
(U.S.A. )Bitter( USSR)
Little Borax(U.S.A.)Big Borax(U.S.A. )Albert(U.S.A. )fJIono(U.S.A. )Na

Mg
moles/m 3

Ca

K

Na

K

Ca

Na + K
Mg + Ca

!'1g

1390

205

1150

350

7

3

1.1

60

2.2

18

1.9

27

10

3.1

14

0.2

3.3

0.7

0

5

3.6

140

1.8

31

7.5

78

4

3.7

260

5.4

35

1.4

48

25

7.3

1570

10.2

175

34

154

5

7.6

165

1.0

20

1.6

165

12

7.7

470

10

53

10

47

5

7.6

165

1.0

17

0.6

165

28

9.4

3640

104

300

6.0

35

50

12

300

2.6

25

0.9

115

27

12

70

2.6

4.5

0.2

27

22

15

100

2.4

4.3

0.2

42

21

23

820

21

31

2.2

39

14

25

150

19

1.0

0.2

8

5

33

270

7.5

1.0

0.1

36

10

250

165

4.5

0.5

0.1

37

5

280

945

29

1.5

0.1

33

15

610

293

500..

•

•

•

,

,

,

I

•

• ••

••

400

••

0..

•

t-

en

--- 300

•

•
••

-.

I

0'

~

E

V

U

,
a

200

-...
(1)

:r:

tfl

too

7.6 ± 0.6)( 10-Scal per atom 3 He

a'

2' 4

6

8

'0

1

1'2

(4

16

CORRECTED TEMPERATURE (OC)
Figure 6-4. Silicate content versus temperature f'or waters from ridge crest hot
springs (upper diagram).
Magnesium content versus silica te content for these same
samples (lower diagram).
These measurements were made by John Edmond and hiscolleagues (642).
Figure 6-5. 3 He concentra tion versus wa ter tempera ture
ho t springs.
These measurements were made by Bill
Oceanographic Institution (276).

for wa ters from ridge crest
Jenkins of the Woods Hole

anomaly. These plots suggest that the waters sampled are mixtures
between local bottom water and a very hot water which has equilibrated with hot basalt.
A lower limit can be placed on the temperature of this hot water from recent measurements made in theso-called "black smoker" vents (intense plumes of very hot water
about a foot in diameter, which shoot many meters up into thewater column precipitating CaS04 and black sulfide minerals along
their paths).
Inside the vents from which these waters emanate,
temperatures as high as 300°C have been measured! From the hightemperature data on the solubility of silicate and on the basis of
the temperature at which the Mg content goes to zero, John Edmondand his coworkers at MIT, who have made these measurements, estimate that the temperature of the hot endmember of this series is
330oC~.
Knowing this, an estimate of the chemical composition of
the endmember hydrothermal water could be made by extrapolating
the concentration versus temperature lines to 330°C:.
In order to estimate the impact of this process on theocean's chemical budget, the total flux of this hot water must be
known.
Bill Jenkins of Woods Hole came up with an ingenious
method to estimate ·this flux.
His measurements of He in samplesof these hot waters reveal that they also contain large excessesof both 3He and of 4 He .
As shown in figure 6-5 when Jenkins
plotted the 3He excess against temperature he obtained a straightline.
The importance of this finding was that the slope of thisline corresponds to what is thought to be the ratio of global h~at
release through ridge crest hydrothermal activity to global
He
release from ridge crests.*
Together John Edmond and Bill Jenkins postulated that ~heratio of the chemical anomalies to either the temperature or He
anomaly in the sample analyzed was typical of the global ratio.
If so, then they had the means to assess the overall importance ofinteraction of sea water on the budget of a gi ven element in 3 thesea.
They- merely mutliplied the chemical anomaly to heat (or He)
anomaly ratio found in the vent samples by the global heat (or3He) release rate.
The estimates obtained in this way are compared with the rate of delivery from rivers in table 6-6.
As can beseen, hydrothermal circulation removes Mg and S from the sea andadds K, Ca and Si to the sea.
The second sink to be considered is the formation of authigenic minerals wi thin marine sediments.
An estimate of the importance of this sink can be obtained from the measurement of thechemical gradients in the pore waters of sediments. For any given
consti tuent the product of the concentration gradient times theeffective molecular diffusivity gives its flux through the particular horizon of the sediment column. This flux provides a measureof the net production or consumption of the constituent beneaththe horizon at which the gradient measurement was made.
Thus bymeasuring the gradients in the concentration of a given ion at*The total heat transport is obtained by subtracting the amount of heatbeing carried by conduction (as measured by thermal probes in the sediment) and the theoretical total heat flow based on models for plate
cooling.
The total helium-3 input is based on the integrated helium-3
excess in deep water divided by the ventilation time for the deep sea.

295

Table 6-6. Comparison hydrothermal fluxes with those from rivers
and to the sediments (all in moles/yr).
These estimates weremade by John Edmond and his coworkers (642).
GAINS

LOSSES
Sediment
Hydro.
Loss
Loss

Constituent

River
Input

Hydro.
Input

Mg++

5.3 x 10 12

-

Ca++

12 x 10 12

3.5 x 10 12Ba++

10 x 10 9

2.4 x 10 9

?

-

Li+

1.4 x 10 10

16 x 10 10

?

-

K+

1.9

X

10 12

1.2 x 10 12

?

-

Rb+

0.4

X

10 9

2.4 x 10 9?

-

4 x 10 12-?

12
4 x 10F-

16 x lOla-?

1 .1 x lOla81

6 x 10 12

3 X 10 12

?

-

S04=

p

~

3 x 10 12<0.1 x lOla

ECO 2

~

2 x 10 13~0.1

x 10 13296

8 x 10 12?

-

3 x 10 10
2

X

10 13

<0.1 x lola-

various depths in the sediment column it should be possible toestimate not only the its flux into or out of the sediment but also the places within the sediment column where these reactions aretaking place.
C~onsiderable evidence has been obtained to showthat pore waters are depleted in Mg++.
An example of the decrease
with depth in the sediment column of the Mg content of pore fluids
is shown in figure 6-6.
The missing Mg appears to replace ironwhich is removed from oxide or silicate minerals to form iron sulfide (a secondary mineral resulting from the conversion of S04=
ion to S= ion by anaerobic bacteria).
Figure 6-7 shows the correlation between S04= and cation loss in pore water samples
squeezed from deep sea drill hole material.
The third sink for cations from the sea is the precipitationof silicate minerals wi thin the water column or at the sedimentwater interface.
Success in determining the extent to which thisprocess contributes has been very limited.
The x-ray diffraction
patterns for recent sediments are dominated by reflections fromdetri tal grains from the continents.
Tha t the K bearing phases
are largely detrital can be shown by the high argon ages obtainedon recent sediment samples (80 to 400 million yearsJ.
Had theseminerals recrl.stallized they' would have lost their 4 Ar. Similarly the 18 0 / 1 0 ra tios and 87 Sri 86 Sr ra tios in these minerals are
more akin to those in continental soils than to those expected forminerals precipitated in the sea. While these observations do notprove that silicate phases are not forming on the sea floor or in
the wa te r column they do make the search qui te difficul t.
Asnearly all of what is seen in x-ray diffraction patterns and inisotopic analysis can be explained by continental detritus there
is no way to pin down the contribution of ocean "precipi tates".
As our knowledge of the removal mechanisms of cations isquite primitive, beyond a few general statements we cannot say
much about the controls acting to maintain these ratios.
One generalization that can be made is that the difference between the
total anion and total cation charge must be balanced by H+ or OHions.
Were such a difference to develop the pH of the solution
would change.
For example, if the ca tion charge were to becomelarger than the anion charge the difference would have to be madeup by OH- ion and the solution would become basic (i. e., its pH
would rise).
The reactions which remove cations to the sea floor
are all increased in rate with a rise in ocean OH- ion concentration. Hence a feedback loop via the ocean's pH would always drive
the ocean' ,s ca tion and anion charge to equali ty.
Any tendency
towards a buildup in cation concentration (i.e., in alkalinity)
would be countered by more rapid cation removal.
Working in concert with the mechanism controlling the cation loss rate would bea feedback loop involving the cation gain rate.
A more basic
ocean would have a lower PCD2..'
Hence continental weathering
rates would also be lower.
While we don't know the specifics ofthese processes the generalization must be valid.
The total
cation charge in the sea must very nearly balance the total anion
charge; otherwise large pH change would occur. As discussed above
the anion charge is fixed largely by the formation and destructionof salt deposits.
The ratio of the ca tions yielding the proper total need not

297

50
E
:r:
I-

a.. 150

w

Cl

200

250 ~_....._-~-...-..._--------------~
10
20
30
40
60
50
o
mol/m 3Figure 6-6. The K+, Mg++, Ca++, and S04 = ion contents of porewater squeezed from various depths in a core taken at 13°N
latitude and 64°W longitude and at a water depth of 1.2 kilometersin the Caribbean Sea as part of the National Science Foundation's
Deep Sea Drilling Program.
These data were obtained by Fred
Sayles and Frank Manheim, Woods Hole Oceanographic Insti tution(529b).

298

65 r - - -.......- - - r - - -......---......----.--......-....
,....,..
rt')

60

E

"-0

E

'-'

55

+
+

•

0

u

+
+
+

50

0'

~

40 .....- .....- -......- -.......--.-AII-------...
o
5
15
20
25
30
(mol/m 3 )Figure 6-7. Plot of the Mg++ + Ca++ versus 804 = ion contents ofpore waters squeezed from sediments for the data shown in figure
6-6 (529b).
The solid line represents the average relationshipderived from these measurements.
The dashed line represents therela tionship expec ted if Ca ++ and Mg++ were the only ions replacing the iron removed from oxide minerals for precipi ta tion by the
reduced sulfur.
The shallower slope of the observed line indica tes tha t ions other than Ca++ and flIg++ are invol ved.
Po tassium
is one candidate, but the decrease in K+ (as shown in figure 6-6)
accounts for only one-third of the slope difference.
The logicalcandidate for the remaining difference is Na+.
Since
the Na+concentration is so large, the loss of the needed amount would not
produce a measurable gradient in the pore water column.

299

of course remain constant wi th time.
Changes in the geometry of
the sea floor produced by plate motions, changes in type of organisms Ii ving in the sea through evolution and changing habi tats,and changes in the oxygen content of the atmosphere will lead torelative removal probabilities for these cations.

POSSIBLE CAUSES FOR PERTURBATIONS
So far we have been speaking largely in the abstract aboutchanges in the operational characteristics of the ocean.
Have wejust been playing thought games or is there a reason to consider
such scenarios?
To answer this question we first need to remind
the reader that during the past million years our planet has undergone very large quasi-cyclic variations in climate called glaciations.
The length of these cycles has been about 100,000 years(hence on the order of the residence time of C: and P in the sea).
It does not stretch the imagination to believe that climaticchanges which built and destroyed ice sheets of continental dimension might have changed the rate of ventilation of the deep sea,
the delivery rate of ions to the sea, and the ecology of the sea.
If so, then these changes must have upset the balance between input and loss and caused the steady state concentrations of thevarious nutrient constituents to shift.
On a longer time scale the movements of the great plates
which make up the earth's outermost shell have surely altered conditions in the oceans for they would have altered the geography ofthe ocean basins and of the continents.
These changes were sufficiently slow that the cycles of the nutrient constituents shouldhave been able to keep pace.
If so, the sea must have undergone
long term changes in the concentrations of these constituents.
Onthese time scales the concentration of other consti tuents mightalso have changed.
For example, the concentrations of 02, Ca++and K+ adjust on time scales of millions of years.
Those of S04=
and Mg++, Cl- and Na+ adjust with time scales of tens to hundredsof millions of years.
In addi tion to the gradual changes associated wi th the longterm al terations of the earth's surface geometry, we know of several geological catastrophes.
The greatest of these occurred 60
million years ago when during a short in terval of time « 10,000
years) over half the species of organisms on the planet became extinct.
Among those to go were all the hard part producing microorganisms ~hich inhabited the Cretaceous surface ocean (694).
There is now good evidence linking these extinctions to the impactof a large asteroid or comet.
Again it does not stretch one's
imagination to believe that the nutrient consti tuent concentrations in the sea must have suffered some perturbations as the result of this major ecologic disruption.
Finally, man is now conducting his greatest geochemical experiment.
CO 2 is being added to the atmosphere at a stupendous
rate by the combustion of fossil fuels.
By the end of the next
century we will have produced an amount of C02 equal to several
times that in the atmosphere prior to man's impacts.
About onequarter of this CO 2 will at that time reside in the ocean.
The

300

remaining three quarters will reside in the atmosphere.
While theimmediate concern to mankind is the climatic warming that this excess atmospheric CO 2 will cause, there will also be impacts ofthis C02 on the chemistry of the ocean, which will last for thousands and even tens of thousands of years.
It behooves us tothink through in advance what the consequences of this "experirnent Tl will be.
Thus the analysis of the control systems we conducted at the
beginning of the chapter were not just a thought game.
These systems must have played an important role in the history of ourplanet and they surely will be brought into action by the perturbations man will induce on the ocean's chemistry.

RECORDERS OF PALEOOCEAN CHEMISTRY
If we are to learn about the changes in ocean chemistry that
occurred in response to glacial to interglacial changes, to geologic ca tas trophes, and to the long term geographic changes induced by plate motions then we must develop a means of readingthese changes from the record in sediments.
This is no simple
task. No one has yet discovered a reliable means by which the concentration of any of the ingredients of sea salt is directly recorded in marine sediments.
We can understand why this is difficult by remembering that everything that enters the sea must
eventually leave.
Thus the composi tion of the ave rage sediment
formed at any gi ven time is related far more closely to the composition of the material entering the sea at that time than to thecomposition of the sea salt at that time.
The chemistry of sea
salt continually adjusts so as to promote throughput of the material arriving from rivers.
These changes need not bear any simple
relationship to the changes in inputs.
By contrasting sediments deposited during the last glacial
period with those deposited since the end of this period (i.e.,during Holocene time) there are several means by which we can gain
information about changes in the ocean's nutrient constituent concentrations.
1).
Through the application of the dating methods describedin chapter 5 it is possible to determine how the accumulation rateof calcite at various places on the sea floor changed between lateglacial and post glacial time.
If these cores are taken at sites
above the lysocline then these measurements gi ve informa tion regarding the, rate of production of calcite by organisms living in
the overlying water column.
Changes between glacial and interglacial time of course could be brought about by a change in ocean
ecology, by a change in ventilation rate, or by a change in the
supply rate of carbon to the oceans.
2).
By studying calcite from cores taken in the transition
layer below the lysocline it is possible to determine how the extent of calcite solution has changed with time at this particularpoint on the sea floor between glacial and interglacial time.
anexample of such a record is shown in figure 6-8.
If such a study
included cores from various water depths, it should be possible todetermine how the lysocline changed between glacial and interglacial time and also how the thickness of the transition zone301

0 3....._ 4....._ ...5_

o

20

40 60

11.000YRS AGO

E

::c
Ia..
w
o
-

126 1000YRS AGO°Fi gu re 6- 8 . CaG 3 an d <5 1 8 0 r e cords for core V19- 2 9 (4 0 S, 84 0 W, 3 • 2km) from the east equatorial Pacific.
This site currently lies
beneath the lysocline for calcite.
The variations in CaC03
content wi th time are mainly the resul t of changes in the extent
of calcite solution (more in warm periods less in cold).
The
CaC03 changes appear to lag the ice volume changes by about 10,000
years. The measurements were made by Nick Shackleton at Cambridge
University (328).

302

(from no calcite solution to complete calcite solution) changed.
Glacial to interglacial variations in the depth of the lysocline
would signal changes in the C03= ion concentration in the deepsea. Changes in the thickness of the transition zone would signal
changes in the calci te rain ra te or in the gradient of CO 3= ion
content in the water column beneath the lysocline.
3). Studies of the relationship between the distribution of
the habitats of various species of benthic foraminifera on the sea
floor and the distribution of dissolved oxygen in the deep sea reveal that certain species compete more favorably at sites bathedin highly oxygenated bottom water while others compete more favorably on sites bathed in poorly oxygenated bottom water.
Thus reconstructions of the distribution during glacial time of speciesof benthic forams with affinity for high oxygen water and of species wi th affini ty for low oxygen water should provide important
information regarding changes in the oxygen content of the deepsea between glacial and interglacial time.
An example of such achange is shown in figure 6-9.
As we have discussed above, suchchanges would signal shifts in the phosphorus content of deep
water.
4). Measurements on samples of Holocene age from borings in
the Greenland and Antarctic ice caps made by Oeschger and his colleagues in Bern, Swi tzerland, suggest that the air trapped insmall bubbles in the ice has a CO 2 content close to that in thepre-1850 atmosphere.
By contrast, samples from glacial time haveC02 to air ratios distinctly lower than those from post glacial
time.
An example of these resul ts will be shown in chapter 9.
Thus a means exists for reconstructing the C02 partial pressure in
the atmosphere (and hence also in the surface ocean) over the last
several hundred thousand years!
5).
As we have seen, the 18 0 /16 0 ratio in sea water changes
between glacial and interglacial time because the water that goes
into building the continental ice masses is depleted by several~~rc1~t inthe heavy isotope of oxygen.
The changes in the0/ 0 ratio of deep sea water produced by the waxing and waningof these ice sheets are recorded by the shells of benthic foraminifera.
As salt is left behind when vapor is carried from the sea
surface, the 180 record kept by benthic forams provides a salinity
record for mean sea water.
Further, as mentioned above, the 180
record allows us to tie together reliably information from deepsea cores taken throughout the world ocean.
An example of such a
record was given in figure 5-12.
6).
There is one co~stituent of sea salt which we have notas yet mentioned; namely L 3C02 .
13 C is a stable isotope of carbon. On the average 1 atom of carbon in the earth out ofis ~f
this type.
As we shall now show, the distribution of
3 C/1 C
ratios in the carbon dissolved in the ocean is closely tied to the
distribution of P0 4 (and N0 3 ) in the sea. Foraminifera record the
13C/ 12C ratio of the carbon in sea water just as they do the
18 0 /16 0 ratio.
By measuring the 13 C/12 C ratios in forams whichgrew at different depths and locations in the ocean at some timein the past we should be able to reconstruct the distributions of
P0 4 (and N0 3 ) for that time!
An example of such a record is shown

9q

303

8 180(%0)

o3

4U.PEREGRINA

(0/0)

o

5

50

100 I

ATLANTIC

OCEAN
44°N

~

E 2

"""'
J:
l-

24°W
3.3 km

3

n.
w 4
0

5
6

Figure 6-9- Plots of 0 18 0 versus depth for the benthic foraminifera species Uvigerina peregrina and of percent Uvigerina peregrina in the benthic foraminifera assemblage versus depth in acore from the flanks of the Mid -Atlantic Ridge at 44°N.
These
results show that Uvigerina peregrina is absent from this part ofthe sea floor during post-glacial time and during parts of thelast interglacial period.
This species is associated wi th highoxygen content waters in today's ocean. The 18 0 measurements weremade by Nick Shackleton at C:ambridge Uni versi ty and the benthic
faunal studies were made by Streeter at the Lamont-Doherty
Geological Observatory (619).

304

8 180 (%0)
45O3......-........
~.................

2.....-..

E

~

.-0-

J:

W
0

4

I

ATLANTIC

OCEAN
25°N
17°W
2.6 km

G

6

I
8

G

Figure 6-10. Plots of 6 18 0 and 6 13 C for benthic foraminifera as afunc tion of depth in a core from the eas tern margin of the North
Atlantic Ocean.
It can be seen that the 6 13 C values for forams
which grew' on the sea floor during glacial time were lower thanfor those which grew during interglacial time. These measurements
were made by Nick Shackleton of Cambridge University (328).

305

in figure 6-10.

THE MARINE GEOCHEMISTRY OF CARBON-13
The ratio of 13C to 12C in the dissolved inorganic carbon in
the sea varies from place to place for two reasons.
First, wh~nplants draw carbon from surface sea water for photos~nthesis, 1 C
is fixed with slight preference over 13 C yielding a 3C/12 C ratio
in plant matter about 20 0 /00 lower than that in the EC02 in thewater in which the plants grow.
When animals consume the plants
they incorporate in their tissues carbon of nearly the same
isotopic composition as in the food they eat.
They also respire
carbon of nearly the same isotopic composition as that in theirfood supply.
Because plant matter is formed only in the ocean
mixed layer, and because some of this material is oxidized only
after falling to deeper layers of the sea, the 13C/ 12 C ratio in
surface water is higher than that in deep waters (i.e., 12C: ispreferentially transferred from surface to deep water through thephotosynthesis-respiration cycle).
The second reason that the 13 C/ 12 C ratio in dissolved inorganic carbon varies from place to place within the sea is that
there is a fractionation between the carbon in atmospheric C02 andthe EC0 2 in surface ocean water. The 13 C/12 C ratio in atmosphericC02 is on the average about 90 /00 lower than that in surface
ocean EC02.
The extent of this fractionation depends on temperature.
The depletion varies from 10.6 0 /00 at OOC to about
7.6 % 0 at 30°C.
Thus the carbon in surface ocean waters fromhigh latitudes should have higher 13 C/12 C ratios than the carbon
in surface waters from low latitudes.
Before proceeding wi th an analyses of how these two sourcesof variation are imprinted on the dis t ribu tion of 13 C/ 12C ratios
in the sea, we must mention the uni ts used by geochemists doing
13C studies.
Like 14C/ 12 C and 18 0/ 16 0 results, 13 C/12 C results
are gi ven _ as the pe r mil difference from that in a reference
sample. The equation is as follows:o13[ 13 c/12 C )sample - 13 C/12 C )refC - 1000
1312C/ C)refJ

6-1

The reference standard is a belemnite from the Peedee formation in
North Carolina used by Urey in his early work (see page 681).
Onthis scale mean ocean carbon is close to zero~ i.e.,o 1 3 Cmoc = 0
The difference between
designated as follows:~o13CA_B =

0/ 00

the

o13C valuesforvarioussamplesiso13 C )A - o13 C )B

Thus the difference between the o 13 C of marine plant tissue andthat of EC02 in the water in which the plant grew is given as:306

~013Cplant_LC02

-20

0/00

A profile of o13 C versus water depth for a station in the
North Pacific is shown in figure 6-11.
Also shown in this figure
is a profile of phosphorus _ The relationship of o13 C to PO 4 intrr~s and other partsof the ocean is summarized in the plot of
o C versus P0 4 concentration shown in figure 6-12 _ There is aninverse correlation between the two parameters. * The reason forthis correlation is that the variations of both parameters in thesea are mainly controlled by the photosynthesis-respiration cycle.
Were this the only source of perturbation then the expected slope
could be approximated by the relationship:
=

~o

13

C/P)organic tissue
Cplant-LC02

LC02 mean ocean water
~o

13

about
Cp l an t-LC02
averagesmarine plants about 105, and the
Thus:
~o13C

=

~P04

-20 x

105
2250

= -0.93

6-2-20 0 /00,
the
C/P
ratio
in
ECO 2 content about 2250 ~m/kg_
0/ 00

~m/kg

Thus the observed and predicted slopes agree reasonably well,suggesting that the photosynthesis-respiration cycle dominates thetemperature dependent air-sea interaction.
If so, then what is the fractionation induced by air-sea
exchan6e?
The o13 C values for the EC02 in surface waters areplotted against water temperature in figure 6-13.
Also shown isthe expected trend for equilibrium with the atmosphere.
Clearlythesurface
ocean
does
not
achievelocal
carbonisotopeequilibrium with the overlying atmosphere.
This is not surprising
for the time required for isotopic equilibrium to be achieved isgiven by the following relationship:[LC02]
1c arbon isotopes[ C0 2 ]

19as

6-3

To exchange the isotopes of carbon in the LC02, the HC03- and
C03= must exchange as well as the C02The ratio of LC02 to C02
in surface water ranges from about 180 in warm water to about 150
in cold water. Thus the carbon isotopic replacement time is about
15 years (and the half replacement time about 10 years).
Surfacewaters simply do not "sit still" this long for local equilibrium*As will be discussed in chapter 10 the addition of fossil fuel andforest soil
~ to the atmosphere during the last 200 years has
reduced the 0 C value for surface ocean water by about 0.80/00.
Hence prior to 1850 the mean 0 C for warm surface water was likely
about 2.7%0 (as opposed to the value of 1.9%0 measured in theearly 1970' s).

cq

307

2
O--.....- ......- - - . . - - - . - - -......---:-T"".~

•

•1•
,.t
.-

,•

••

• •• •

....
..••
.--•

13

~

I
I-

a..
w
3
o

4

C

••
•
~BOTTOM

5 O'-----'-----~2---~3~......

P0 4 (,umol/kg)
Figure 6-11. Plots of 0 1 3 C and P0 4 versus depth at GEOSEC.S stat~~n346 (28°N, 121 0 W) in the northwestern Pacific Ocean.
TheC
analyses were made by Craig and Kroopnick at the Scripps Ins titution of Oceanography (292) and the P0 4 analyses as part of the
GEOSECS program (425).

308

,
\

\

J~ WARM SURFACEWATERS (CORRECTED)

2

EX PECTED SLOPE= fj, al3cPHOTOx %)ORli[1: C02] MEAN SEA;20 x (05 =-0..93~
WARM

SURFACE
WATER
EARLY

2250
08S.SLOPE=-O.93

0/00}J-mol/kg
0/00

fLmo

I/k
9

DEEP
ANTARCTIC

19705(MEAN 1.9%0)ao3Figure 6-12.
Relationship between o13 C and P0 4 in the world
ocean.
The o13 C values for warm ocean water have been correctedby o. BO 100 to remove the decrease which has taken place over thelast century as the result of the invasion of CD 2 produced by theburning of fossil fuels and the decline of forest-soil reservoirs.
The slope defined by these results is close to that
expectedfromphotosynthesis
andrespiration
cycle(i.e.,-O.93 0 /00/ o C).
The 13 C data on which this diagram is based wereobtained in the laboratories of Harmon Craig at the Scripps Institution of Oceanography and of Peter Kroopnick at the Universityof Ha wa i i ( 29 2 , 29 3 , 29 4, 29 5 , 296 , 29 8, 30 1) •309

EQUILIBRIUM:3

WITH ATM

• ••

-° 2
-°
~

U

-GO
rt')

I

~
~

• •

4t

8

88

0
/ 00

(J

(J

8

,_ eel':•al3 c = -7.2

o

0

O0j!°
'&0
.°
00.
auo
8

0·

1m
til

8

o

a

0

e
NORTHERN•ATLANT'C
PACIFIC•CENTRAL<:)ANTARCTIC

8~00..-......- - - - -.......- - - - -.......- - - - -...
o
10
20
30
TEMPERATURE (OC)

Figure 6-13. Plot of 0 1 3 C measured on the ECO 2 extracted fromsurface ocean wa ter as a func tion of surface wa ter tempera ture.
The straight line shows the
013 C forequilibrium with theatmosphere at the time these samples were collected.
Clearly thesurface wa ter resul ts do not follow the trend expected if local
isotope equilibriwn were es tablished wi th the overlying a tmosphere.
The 13 C data are from the references cited in the figure
6-12 caption.
The equilibrium values are calcula ted from da ta
obtained by Mook of Groningen in the Netherlands (297).

310

to be achieved.
They are replaced by vertical or horizontaltransport on the time scale of a few years.
So we see that local
1~chtgge with the atmosphere does not make a strong imprint on the
C/ C ratio distribution in the sea.
l\he 1 tong exchange time for carbon is otopes also influences
theC/ C distribution in the sea.
We have alread~ seen thatnO~here did surface ocean water have the atmos~heric 1 C/ 12 C ratio
prlor to the onset of nuclear testing.
The 1 C deficiency ranged
from 4-5 percent in the temperate surface ocean to more than 10
percent in the surface waters of the Antarctic.
We shall havemore to say about the implications of this slow exchange in thechapters to come.
The iI¥,:0rtant point to emphasize here is that in today' s
ocean 13 C/1 C ratio differences from water mass to water mass provide an excellent measure of PO~ differences between these masses.
If the temporal record of 13 C/1£C ratio in the shells of foraminifera which grew on the sea floor (benthic forams) and in theshells of foraminifera which Ii ved in surface water (planktonic
forams) provide a proxy record of the temporal variations in theo13 C for these waters between glacial and interglacial time, wehave the wherewithal to reconstruct the changes in the P04 content
of various water masses between glacial and interglacial time. As
we shall see in chapter 9 this will prove very important to reconstructions of the chemistry of the glacial ocean.
THE URANIUM CONTENT OF CORAL

Let us consider one last paleoocean indicator which may give
us information about changes which occurred on the time scale of
10 million years.
Uranium is present in the sea in remarkably
great quantities compared to its abundance in crustal rocks. It isa biounlimited element with a reactivity so low that it remains inthe sea for about 700,000 years before being removed to sediments.
Measurements on saline lakes also reveal high uranium contents.
In fact, the U content of the waters of these lakes is
proportional to their total dissolved carbon content. Pyramid Lake
in Nevada, for example, has a tenfold higher EC02 and a tenfold
higher U content than sea water. Mono Lake in California has ahundredfold higher ~C02 .and a hundredfold higher U content.
We
can imply from this that if, in the past, the ocean had a different EC.o2 concentration it probably had a different U concentration, too.
Uranium owes its remarkably low reactivity in natural
waters to' the fact that it forms highly stable complexes with
C03= ion.
Uranium has another peculiarity.
Corals incorporate uranium
in their CaCO 3 in nearly the same proportion to calcium as foundin sea water.
Apparently, these organisms cannot distinguish between the doubly charged cations of ITa, Sr, Ba, Ra, and U, and use
them indiscriminately in building their CaC03 reefs. Thus the U/Ca
ratio in a well-preserved marine coral should be an index of the
U/Ca ratio in ancient sea water. More important, since the U content of the sea apparently reflects its e03= ion content, if
corals grown at some time in the past are found to have a different U/ea ratio, then the water in which they formed probably had a311

different C03=/Ca++ ratio than today's sea.
Figure 6-14 shows histograms of the U contents of corals forfour time intervals:(a) the last 8000 years, (b) 80 000-220 000
years ago, (c) 2-5 million years ago, and (d) 5-20 million y~arsago. No significant differences are found back to 5 million years.
The U/Ca ratios in corals from each interval in this time period
are much the same as the U/CE. ratios in the living corals in today's ocean.
In the 5-20 million year interval, however, the U/Ca
ratio ~s an average of only half as great as it is in living corals. Slnce the corals used for this study show no conversion ofaragoni te to calci te and since they yield uranium-helium ages *consistent wi th those based on the geologic ages of the deposi ts
in which the corals were found, the low U content of the oldest
group of corals does not appear to be the result of U loss during
the millions of years which have elapsed since these corals formed.
From studies of marine sediment borings we know that during
this 5-20 million year time period the saturation horizon separating CaC0 3-rich and CaC0 3-poor sediments was at approximately thesame elevation as it is today (see figure 2-31).
We can conclude
then that the product of the Ca++ and C03= ion contents was alsoroughly the same as it is today.
If this is true, to have had a
twofold lower C03==/Ca++ ratio, the Ca++ ion content must havebeen 1.4 times tha t of today and the CO 3 == ion content mus t havebeen 0.7 times that of today. The uranium content would then havebeen 0.7 times that of today and the U/Ca ratio one-half thatfound in today's ocean.
If so then the Ca content of sea water must have decreased
during the last 20 million years.
If the anion content (Cl- +S04 =) has remained the same, then the Ca++ decrease must havebeen accompanied by an increase in Mg++, K+, or Na+.

FACTORS

INFLUENC~NG

NUTRIENT GRADIENTS IN THE .DEEP SEA

We have seen that the present mode of ventilation of the deepsea (i.e., from northern Atlantic through Antarctic to the Indianand the Pacific) has a strong impact on the distribution of properties in the deep sea.
The concentrations of the nutrient constituents (phosphate, nitrate, silicate, ~C02, alkalinity and barium) all increase along the path of this flow.
That?f dfssolved
oxygen and of carbonate ion decrease.
The ratios of 4 C/ 2c. (because of radiodecay) and that of 13 C/ 12 C (because of respiration)
also decrease. This pattern is imprinted in deep sea sediments inthree ways:
1)
The depth of the lysocline decreases along the path ofthe deep current.
2)
The 13 C/ 12 C ratio in benthic forams decreases along thepath of the deep current.
3) The abundances of those species of benthic forams adopted
to low oxygen water increase along the path of thecurrent.
By studying the change in these parameters with time in cores*The uranium-helium method is closely analogous to the potassiumargon method we discussed in chapter S.

312

40

20
LIVING HOLOCENE'

(f)

w

,I

-'

Cl..

~

5

(f)

0

PLEISTOCENE

I
I
I

15

<l
lL

0

0
0::
W

6

2 (0) 3

4

5

0

0

2 (b) 34

52 Cd) 34510

PLIOCENE

CD
~

::>

4

z

5
2

2 (e) 3

4

5

U CONTENT (ppm)

Figure 6-14. Histograms showing the uranium contents of corals in
four
different age groups.
In each case, the U content expected
if the corals formed with the same U/Ca ratio found in today's
ocean is indicated by a dashed line.
The data are shown in (a)for living corals and those formed during the last 8000 years (thepresent interglacial period); in (b) for corals which lived during
the two previous periods of interglaciation (80,000-220,000 years
in age); in (c) for corals from the Pliocene Epoch (2-5 million
years in age); and in (d) for corals from the Miocene Epoch (5-20
million years in age).
These uranium measurements were drawn froma variety of sources listed in reference (547).

313

taken throughout the world ocean, it should be possible to learn
about changes in the sources and flow-patterns of deep water.
We
will have more to say about this in chapter 9.
SUMMARY

In this chapter we have seen that there is good reason tosuspect that the concentrations of the various nutrient constituents in the sea as a whole, and even in one part of the sea relative to another, have changed with time.
They have been drifting
about in an effort to minimize the differences between loss andsupply created by geologic and climatic changes.
The glacial cycles of the recent past provide the ideal opportuni ty to searchfor such changes.
As the length of these cycles is similar to theresidence time of a number of the key constituents, chemicalchanges are likely to have occurred.
We have complete and welldated records of deposition over the last complete glacial cyclefrom many places on the sea floor.
Much of chapter 9 will be devoted to an attempt to piece together and interpret this record.
Man's activities will alter the course of ocean chemistry.
The CO 2 produced by burning of coal, oil and natural gas is thechief culprit in this regard.
In
chapter 10 we will study theinterplay between man's C02 inputs and nature's control systems.
The answer to the question,"What Keeps the System in
Whack?", is negati ve feedback controls. If the throughput of anyof the constituents of sea salt gets out of balance, its concentration will begin to change.
The change in concentration will
feed back by altering the rates of formation of the particulate
phases which carry the constituent to the sea floor and by altering the degree of preservation of these phases on the sea floor.
These feedbacks tend to eliminate the imbalances created by geologic and climatic change.
It must be emphasized that we have asof now only a slight glimmering regarding the identity of the important control mechanisms.
Thus we are bound to be in for surprises. The main point to be made here is that the ocean's chemistry has almost certainly changed wi th time and that economic
logic is an important ingredient in the understanding of thesechanges.
PROBLEMS

6-1

Measurements on an equatorial Pacific Ocean core reveal thefollowing accumulation rates for post glacial time:
Organics<0.01
g/cm 2 10 3 yr
Calcite1.0
g/cm 2 10 3 yr
Opal
0.2
g/cm 2 10 3 yr
Other
0.5
g/cm 2 10 3 yr
Total
1.7
g/cm 2 10 3 yr
Sediment traps from the same area yield the following rain
rates:
Organics
10
g/m 2 yr
Calcite10
g/m 2 yrOpal
20
g/m 2 YrOther
5
g/m 2 yrTotal
45
g/m 2 yr314

Assuming that the rate limiting step for opal dissolution is
the pore water resaturation, what must S*opal be?
In the same core measurements for late glacial time
reveal the following accumulation rates:
Organics<0.01
g/cm 2 10 3 yr
Calcite1.5
g/cm 2 l0 3 yr
Opal
0.7
g/cm 2 10 3 yr
Other
0.6
g/cm 2 10 3 yr
Total
2.8
g/cm 2 10 3 yr
If tQe silicate concentration in deep sea water during lateglacial time was the same as it is today, what must have been
the rain rate of opal during glacial time?
6-2

Assume that the input of silicate from rivers to the two-box
ocean varies with climate.
Assume further that the climatic
cycles are rectangular with each glacial and interglacial interval 40,000 years in duration.
If the river input of silicate is twice as high during glacial times as during interglacial times, what is the temporal pattern of the silicate
concentration in the deep sea reservoir? Assume that all thesilica reaching the surface is utilized by plants.
Assumethat in interglacials 97% of the opal is recycled and that inglacials 99% is recycled.
Assume that the mixing rate in thesea is twice as great during glacials than during interglacials.

6-3

Assume that during glacial times the mean C/P ratio in organic matter falling from the surface to the deep sea is 115
while that during interglacial time is 95.
Assuming that
this change occurred abruptly at the end of glacial time how
would the O2 content of deep water change over the first few
thousand years of interglacial time?
How would the difference between the o13 C for surface water and for deep watercarbon change on this time scale?
Would the o13 C change belarger for deep water or for surface water?
Why?
Assumethat the NIP ratio in marine organic residues remains cons tant from glacial to interglacial time.
No changes in thechemistry of rivers or sea occur nor do any changes incirculation pattern or rate.

6-4

If 50 million years ago the Ca content of the sea was fourtimes today' s value and if at that time the lysocline stoodat today's mean depth, then according to the hypothesis presented in the text, what would have been the uranium content
of corals living at that time?

6-5

Thick salt deposits found beneath the floor in the Mediterranean Sea have an age of 4.5 million years.
They cover an
area of 2xlO 11 m2 and have an average thickness of 1000m.
They consist of about half CaS04 and about half NaCI. By how
much would the deposition of this salt have changed theocean's salini ty?
By how much would it have changed its Ca
content?

315

SUPERPROBLEM#6

Jim Brie, a marine geologist interested in the earth's climatic record, puzzles over the influence of bioturbation and changing sedimentation rate on the character of the record contained indeep sea sediments.
To study these effects he constructs a simple
sediment machine which reproduces what he considers to be themajor elements of the problem.
His machine consists of a mixing
hopper and an accumulation column.
Batches of new sediment areadded to the hopper, thoroughly mixed with what is already there
and then batches of the same size as those added are dropped fromthe hopper onto the sediment column.
After each batch is dropped
the hopper is raised an appropriate distance before the procedure
is repeated.
Each batch of sediment represents 1000 years of sediment rain.
Jim Brie approximates glacial to interglacial cycles
by varying the rain rate of input materials.
He assumes that individual glacial and interglacial periods have a duration of
40,000 years and tha t the transi tions from one clima te to theother are quite sharp.
For his first experiment he adopts the following batch sizes:
Glay
GaGa 3Benthic
Ptanktonic
80 0 (Planktonic)1
o 8 0 (Benthic)
23 0Th Atoms

GLAG.IAL
400 grams
800 grams
2,000 shells
10,000 shells
2.0 0 /00
-2.0 0 /00
10 10

INTERG LAG~IAL
200 grams
600 grams
10,000 shells
1,000 shells
0.0 0 /00
-4.0 0 /00
10 10

Prior to the addition of each batch the hopper contains 4,000
grams of sediment.
If Jim Brie's sediment column has a cross-sectional area of
250 cm 2 and if the clay makes a sediment of mineral densit~ o.6g/
cm 3 and the GaG0 3 a sediment of mineral density 1.0 g/cm, what
are the depth profiles of percent GaGa 3, of the ratio of benthic
to planktonic forams, of the 0 18 0 for the planktonic shells and of
the 0 18 0 for the benthic shells?
Jim Brie sends the profiles he obtains in this way to afriend who has the bad habit of reading sedimentary records in arather simplistic way.
If this friend were to assume that age increases linearly with depth and were he to disregard bioturbation
effects, wnat faul ty conclus ions might he come to with regard to
the shape of the climatic cycle and to the phasing of the various
changes?
In his calculations Jim Brie first treats 230 Th atoms as
stable.
Having obtained the 230Th profile he then corrects foraging (using time elapsed since each batch left the hopper plus
the time of residence in the hopper).
He then plots the logarithmof the ratio of 230Th atoms to sediment mass against depth in thecolumn.
Again he sends these results to his friend.
What false
conclusions might this friend come to with regard to the sedimentation rates for this sediment?

316

Ckapte/? 7
FREIGHT TRAINS AND FICKIAN CONFUSION
THE MOVEMENT OF WATER THROUGH THE DEEP SEA

INTRODUCTION

In this chapter we will take a closer look at what our
tracers have to say abou t the ventila tion of the deep sea.
We
have already established a rough time scale for the renewal ofthese waters and have seen something of their pattern of flow
through the deep sea.
What we seek is a more realistic quantitative representation for this process.
As we shall see thisremains one of the major unsol ved problems in oceanography.
Thematerial in this chapter constitutes a status report on progress
towarti this difficult goal.
TYPES OF MOTION

One of the problems lies in the definition of the nature of
the motions of the wa ter wi thin the deep sea.
In chapter 1 we
spoke of this flow as if it were a river of water moving from thesource regions toward the remote regions of the deep sea.
Whilesuch a flow is the net result of the combined motions in the deepsea, were we to follow any single wa ter molecule we would findthat its pa th was highly tortuous and it would be found moving
upstream and across stream almost as frequently as downstream.
Inthe course of its trip through the deep sea it would log 100 to1000 times more kilometers of movement than were it to have simply
followed one of the vectors shown in figure 1-12.
In an attempt
to simplify the motions in the ocean, oceanographers often separate them into two categories; advection and diffusion.
Advection
constitutes net movement along a major current.
Diffusion is therandom mixing accompl ished by the hos ts of edd ies am backwashes
which are everywhere at work in the sea.
A comparison might bemade to flies in a freight train.
As the train moves along the
track,
the flies move chaotically about in all directions.
Mathematically, advection is described by a velocity (dim~nsions
em/sec) and diffusion by mixing coefficient (dimensions em /sec).
The latter was introduced early in the 19th century by a chemist
named Fick as a way to describe mixing on a molecular scale.
In addi tion to the horizontal motions which carry wa ter fromone geographic locale to another there are also vertical motions.
We have already discussed the gradual upwelling of the water
spreading through the deep sea.
It advects toward the surface.
Again, this motion is not unidirectional.
Superimposed on the317upwelling motion are mechanisms which stir the rising water backdown the column.
Thus, the phys ical oceanographer is faced wi th the rather
awesome task of describing this complex of vertical and horizontal
motions.
His main tools, current meters and neutrally buoyant
floats allow him to study the chaotic motions which we classify
here as the diffusion part of the horizontal motions.
However,
unless placed in very special locations on the sea floor wherestrong unidirectional currents flow, these devices tell us little
about the advective motions which carry water from one ocean basin
to another.
Only in special cases can they tell anything about
vertical motions.
To get at the horizontal advective transports,
the physical oceanographer uses a theoretical argument which relates the flow of water to the tilt of horizons of constant pressure.
Only with the aid of yet to be implemented satellite topographic techniques can this so-called geostrophic technique give
absolute rates of flow.
The physical oceanographer has no way to
directly measure vertical advection or vertical diffusion.
Because of the inability to get at the major motions of interest in connection wi th deep sea ventilation by physical measurements, the tracers we have been describing are looked to as thegreat hope for working out the roles of "freight train motion" andthe roles of "Fickian confusion" in the ventilation of the deepsea.
As we shall see, while the distributions of these tracers
will add much to our knowledge, they will not by themselves provide the solution to the problem.

ONE DIMENSIONAL ADVEC~ION - DIFFUSION MODEL
An example of advection and diffusion in a vertical column
will help the reader comprehend the interaction between these two
types of motion.
To make the example as much like the ocean as
possible, ·we will make our column 4000 meters high.
Sea water
with a temperature of 2°C: and a salinity of 35.0 0 /00 is pumped
into the bottom of the column so that an upwelling velocity of 4
rn/yr is genera ted. *In other words, the volume of water pumped
into the base of the column (in m3 ) in a year is equal to 4 meters
times the cross-sectional area of the column (in m 2 ).
At the topof the column there is a well-mixed layer 100 meters thick (stirred with a mechanical mixing device).
Interactions with the overlying atmosphere maintain this layer at a temperature of 25°C..
An
excess of . evaporation over rainfall maintains the salinity at
36.0 0 /00.
In the absence of any mechanism for vertical stirring
beneath the 100-meter-thick surface layer, the temperature andsalini ty would be the same as in the input water all the way tothe base of the mixed layer.
Here they would change abruptly to*Because the 14C- b ase d ventI. lation'.
tIme
for the deep sea is about1000 years, oceanographers envision an upwelling at theabou t 4 me ters per year (t he mean depth of the ocean
meters).

318

rate ofis 3800

the surface values. *Water would of course have to be pumped
off the top of the column at the same rate that it was added at
the bot tom.
This wa ter would be cooled and reduced in salini ty
before it was recycled.
(In the real ocean the water would flow
along the surface toward the poles where it would be cooled bycontact wi th the atmosphere and reduced in salini ty through thereceipt of an excess of rainfall (and continental runoff) overevaporation.
A more realis tic tempe ra ture and salini ty profile would beachieved if diffusive mixing within the column were included.
Asshown in figure 7-1, such mixing carries the warm and saline surface water downward into the upwelling ~lume.
The characteristic
penetration depth of the surface water z would be given by:
Z

* =-K

w

7-1

and the half-penetration depth,
zl/2,
is equal t~ 0.693z.*
Where K is the coefficient of diffusive mixing (units m /yr) and w
is the upwelling velocity (m/yr). We have already seen that for K= 0 (no mixing), zl/2 is zero (no penetration beneath the mixed
layer).
To produce a half-penetration depth of 400 meters would
require a K of 2300 m 2 /yr (or 0.73 cm 2 /sec).
Temperature profiles
from the Pacific Ocean show a penetration of warm water to a characteristic depths up to about 1400 meters.
The K needed to
producethispenetration
is
5700 m2 /yr(or
1.8
cm 2 /sec) .
Temperature profiles produced in our column by various choices forthese values of wand K are shown in figure 7-1.
The distribution of radiocarbon can be used in quite anotherway in connection with our simple 1-D ocean.
Let us assume thatthe [).14 C. for the surface m'ixed layer carbon is at steady state-50 0 /00 and that the [).14C~ for the water pumped into the bottomof the col-umn is -160 0 /00 (in the Pacific Ocean, after which ourcolumn is patterned, the new bottom water entering from the Antarctic has this [).14 C: value).
We will also assume for the sake ofclari ty that the column is abiotic (i. e., that the carbon moves
only with the water rather than in part sinking down the column inparticulate form).
As 14C. undergoes radiodecay within the watercolumn, the 14 C/ C ratio will be lower in the interior of the column than at the ends!
The extent and vertical distribution ofthis deficiency will depend on the rates of upwelling and diffusion.
Examples are shown in figure 7-2.
Thus, had we come uponsuch a column, and wished to determine its operating characteristics, we would need only to measure the depth profile of salinity(or potential temperature) and the profile of 14 C/ C ratio.
The
former would yield the ratio of K to wand the latter would permit
the absolute value of w (and hence also of K) to be established.
Were this 1-D model to apply to the real ocean, it would permit the utilization rate of oxygen to be determined.
This can be*To maintain the temperature in the surface layer a flow of 10 42calories/ cm yr from atmosphere to sea would be necessary (i.e.,
to balance the cooling caused by the inflow of the cold water frombelow).
To maintain the salinity of the surface water a net
evaporation rate of about 11 cm/yr would be required.

319

SAL IN ITY(%o)

POT. TE M P. (OC)

o o

10

20

30

K=0.73
1.L=0.4..........

E 2~......,
J:

2

K=I.8 C~

I-

a..
w 3

35

36

K=0.73
1.L =0.4
2""2

K=I.8C~2

1~=I.Okm

1~=1.0km

04

W =4.!!!.
yr

5 --................-...._.........a.-. .
Figure 7-1. Plo ts of potential tempera ture and salini ty versus
depth in the hypothetical column discussed in the text.
In eachcase the upwelling velocity is 4 meters/year.
The situations for2
the eddy mixing coefficient e~ual to 0.0, 0.73 and 1.8 cm /sec are
shown.
The corresponding z
values are 0.0, 0.58 and 1. 44 kIn
and the corresponding zl/2 values are 0.0, 0.4 and 1.0 krn.
As
can be seen for the largest of these diffusivities, the warm andsal ty wa ter penetra tes all the way to the bottom of' the column.

320

~14C (0/00)~ 14C (%0)-500-400 -300 -200 -100

o -500 -400 -300 -200 -100

o

-2
E

-

~

:I:
~

a..
w 3
o
4

5....._ ....- -....---...........

_A-_~

Figure 7-2. Plots of 14 C/ C ratio (expressed as ~14C) as a functionof depth in the hypothetical column discussed in the text forvarious upwelling velocities.
Cases for a half penetration depthof o. 4 kIn are given on the left and those for a half penetra tion
depth of 1.D km on the right.
As would be expected the larger the
upwelling ra te the smaller the mid depth minimum in ~ 14 C•
In allcases the ~ 14C for the water entering the bottom of the tank istaken to be -160 0 /00.

321

seen if we allow our column to be inhabited by organisms. We willassume that the water entering the bottom of the column carries
with it 235 llm/kg of dissolved oxygen and that the water at the
top of the column carries an oxygen content of 205 llm/kg. We willfurther assume that the rate of oxygen utilization is the same atall depths beneath the mixed layer.
The profiles of oxygen forvarious va·lues of zl/2 and oxygen consumption rate are shown infigure 7-3.
Were only the ocean functioning in accord with the assumptions of the I-D model!
The distribution of tracers could bereadily harnessed to yield all sorts of useful information.
Unfortunately it does not.
The rates of upwelling vary with 10ca tion and depth (they may even go negati ve in some regions anddepth intervals).
Similarly, the rate of vertical mixing varies
throughout the interior of the sea.
Furthermore, as gradients inthe properties of tracers are found in the horizontal as well as
in the vertical, horizontal advection and diffusion will influence
the vertical gradients of the properties at any given place in thesea. As we have implied from the distributions of 230 Th and 231 pa
in sediment traps, a molecule can be mixed basin-wide along its
own densi ty horizon in a time which is qui te short compared tothat for its transit up the water column.
One more concept has to
be introduced before the reasons for the inadequacy of the I-D
approach can be fully appreciated.
We shall call this the conceptof "outcropping isopycnals".
As stated in an earlier chapter anisopycnal surface is one along which potential density remains
constant.
The movement of water along such a surface requires nowork against gravity.
Thus these are the preferred surfaces along
which lateral mixing takes place.
In the interior of the ocean
basins these horizons are sufficiently close to horizontal thatthe distinction between horizontal and isopycnal mixing is notimportant.
At high latitudes, however, these horizons rise toward
the surface. Thus, instead of speaking of horizontal and verticalmixing, it is more appropriate to speak of isopycnal and diapycnal
mixing (by diapycnal we mean perpendicular to isopycnal surfaces).
Equally important to the changes in the geometry of mixing
which comes about when isopycnals are adopted as the preferredhorizons for mixing is the fact that some isopycnal surfaces "outcrop" at the sea surface. The existence of these outcrops creates
two new mechanisms for the transport of properties across isopycnal horizons. The first is that wind driven currents operating at
the ocean surface will carry water across isopycnal surfaces andin so doing produce diapycnal mixing.
Also seasonal cycles intemperature and salinity will move the isopycnal horizons along
the ocean surface!
The second is that gases can leave the ocean
at the outcrop of one isopycnal horizon and reenter the ocean at
the outcrop of another.
Since 14 C has the gaseous form 14 CD2 ,this latter mechanism becomes important when we use the distribution of natural radiocarbon to calibrate models for the transportof other oceanic properties.
We must conclude at this point that while the two-box and theI-D advection-diffusion models constitute important preliminary
steps in understanding what goes on in the ocean, neither isadequate.
Oceanographers are facing up to the difficult task of322

O2 (fL rna II kg)

o

-~

-

100

140

180

02 (fLmal/kg)

220

100

140

180

220

2

:c

I--

a..
o

w 3
4

11=O.4km

112 =I.Okm

m
W=5yr
2

5 ....----..---.--......~~~m
W=5 yrFigure 7-3. Plots of dissolved oxygen content versus depth for two
sets of dynamic conditions.
For both the upwelling rate is 5m/yr.
For the example on the left the half penetration depth of
the surfac~ temperature and salinity anomalies is 0.4 km.
For theexample on the right it is 1.0 km.
In each example curves forthree oxygen consumption rates are shown (i.e. for 0, 100, and 200].1ffi/ R, 10 3 Y r s ) •323

creating a truly three-dimensional model.
Such a model, however,is a long way off.
Right now we can only examine some of its
pieces.

TRACERS FOR DIAPYCNAL AND ISOPYCNAL MIXINGOne of the pieces of the great 3-D ocean puzzle is the matrixof values for the coefficients of isopycnal and diapycnal mixing.
As these parameters almost certainly vary from place to place andfrom depth to depth within the sea, their dependence on the measurable properties of the sea (i.e., density gradient, current
shear ... ) must be determined.
Three tracers are particularly useful in this regard.
One ofthese is the isotope 2 2Rn , whose application to establishing therate of gas exchange across the air-sea interface we have alreadydiscussed.
Radon gas is produced in sea floor sediments in large
amounts compared to its production in sea water.
Much qf this6
radon escapes the mineral grains upon which its parent 2 Ra isabsorbed and takes up residence in the water-filled pore space between the grains.
During its mean lifetime of 5.5 days much of
the radon occupying the pores of the upper centimeter or so of thesediment escaQes to the overlying bottom water, producing an
excess in the 222 Rn activity in these waters (over that maintainedby the 226 Ra dissolved in these waters).
The distribution of thisexcess above the sea floor carries information about the rate ofvertical mixing of the sea.
The second tracer is 228 Ra with a mean lifetime of about 8
years.
This trace r is formed from the decay of 23 2Th .
As 23 2Th
has such a low abundance in sea water the in situ produced 228Ra
is not meas urable.
Like 22 2Rn , 22 BRa is Rroduced in copious
amounts in deep sea sediment.
Some of the 22 Ra generated in thesediments diffuses through the sediment Rore waters into the overlying sea,- and so the distribution of 228 Ra in the abyssal sea
also has something to tell us abou t ra tes of mixing in this domain.
There is one major difference between these two tra~~~s.
The half-life of 228 Ra (5.8 years) is 550 times that ofRn(0.0105 years).
As the distance a radio isotope can travel bydiffusion before undergoing decay depends on the square root oftime, 228 Ra can travel about 24 times further than can 222Rn !
The third tracer, 3He , is released at the cres ts of the midocean ridges.
These linear features act as would ~ long line ofchimneys releasing smoke to the atmosphere.
This He is carriedaway from 'the ridge crest by the prevailing oceanic "wind".
Itshould also be spread in the vertical by any diapycnal mixing processes at work in the deep sea.
Unlike the isotopes 222 Rn and
228Ra, the excess 3He in the sea does not disappear from the waterby radiodecay.
Instead, it makes its way to the sea surface andfrom there escapes to the atmosphere.
Rather than giving us thekind of absolute mixing rate information potentially availa~lefrom the distribution of 228 Ra and 222 Rn , the distribution of He
has something to tell us about the relative magnitudes of isopycnal and diapycnal mixing in the deep sea.

324

MIXING RATES BASED ON THE RADON-222 AND RADIUM-228

As 222 Rn and 228 Ra have so many things in common it is appropriate to treat their distributions together.
If we make two
major assumptions we can construct a rather simple model withwhich to interpret their distributions.
The firs t assumption istheir distributions in the vertical are established entirely bydiapycnal processes.
If so then the model can be one-dimensional.
The second is that on these short time scales the role ofvertical advection is small compared to that of vertical diffusion.
If so, the I-D model need involve only the balance between diffusion and radiodecay.
As we shall see, while the second ofthese assumptions rests on a firm foundation, the first is shaky.
If the excess 222Rn and 228 Ra given off by the sediments arecarried away from the sediment mainly by vertical mixing then it
can be easily shown that their vertical profiles are governed by
the following equation:
Cz = Coe- fA zj~

7-2

where Cz and Co are, respectively, the concentrations of theexcess at some height z above the bottom and at the sediment waterinterface, K is the diffusion coefficient and A is the decay
constant (i.e., O.693/tl/2) for the isotope.
This equation
calls for an exponential drop off in the excess wi th distanceabove the sea floor.
The distance zl/2 over which the concentration decreases by a factor of two is related to the vertical
diffusivity and decay constant by the following equation:z 1/ 2=o. 693 JfK_~~

7-3

or since we are interested in using the distribution of the excess
to determine K:
2
ZI/2
K

=

[ 0.693]A

7-4

Two examples of sites where the profiles of both excess 228 Ra
and of excess 222Rn have been measured are shown in figure 7-4.
In each case a plot of potential temperature versus depth is also
given.
These plots provide some very important clues as to what
determines the rate of vertical mixing in the deep sea.
First ofall, as expected, the 228 Ra excess extends at both sites muchfurther above the bottom than the 222 Rn excess.
Next the 222 Rn
profiles show the expected exponential shape.
This is best shownby plotting the logarithm of the excess radon content against distance off the bottom.
If the equation given above is obeyed, thenthe points should fall along a straight line on such a plot.
Asshown by the inserts in figure 7-4 this is the case.
Second, theshape of the 228 Ra profile shows a marked change at about the325

POTENTI AL TEMP. (OC)1.50

POTENTIAL DENSITY(%o)

1.55

5.4

1.60

5.5

=
E

5.6
012
LOG £XC. Rn

I

~

~ 5.7

o

5.8
5.9 0

10

20

30

40

34.875EXCESS 222 Rn ACT I V I T Y(dpm/IOOkg)

34.90

SALINITY (%0)

POTENTIAL TEMP. (OC)

0.60
5.3

0.65

0.70
u

!- 1.0

5.4

a.:
:E

=
E

.... -... ..... _, - ~-' 0.9 -.22 %0
oC
«
STN
.'
i= 0.8
z
263

5.5

~5

I

~

Q.

~ 5.6

W

o

o

K>70c~

5.7
5.8~-..I-_......I-_---&_---I

o

25

50
75
100
EXCESS 222 Rn ACTIVITY(dpm/lOO kg)

2

~.

w

•

6 0 .7

•

a.

6

.0

0.6~-----£-.......&._~.-I

.1

.2

.3

.4

EXCESS 228 Ra ACTIVITY(dpm/IOOkg).5

34.67534.70

34.675SALINITY (%0)

Figure 7-4. Plots of excess 222 Rn and 228 Ra activity versus depth
at two GEOSECS stations (left hand and middle diagrams respectively).
The upper panel shows the results for station 32 (24°N,
54°W) in the Atlantic.
The lower panel shows results for station
263 (17°S, 167°W) in the Pacific. The potential temperature pro. files (open circles) are shown along wi th the isotope profiles(solid circles).
Also shown are plots of salinity (open circles)
and potential densi ty (c losed circles) versus potential temperature over the depth range in which excess 22 8 Ra acti vi ty is found(right hand diagrams).
The density gradient for the depth interval over which excess radon is found is calculated by multiplying
the potential temperature gradient in that zone by the ratio ofpotential densi ty change to potential temperature change in thezone 0.5 to 1.0 kilometers above the bottom.
This gradient cannot
be established in the lower half kilometer of the water column
because the salt gradient is too small be to measured.
The hydrographic and radon data were obtained as part of the GEOSECS
program (81).
The 228 Ra measurements at station 32 were made at
the Lamont-Doherty Geological Observatory and those at station 263
at the University of South Carolina (91).

326depth where the potential temperature gradient (and hence density
gradient) shows a marked steeirening. Below this break the density
gradient is small and the 22 Ra content is nearly uniform.
Above
this break, the densi ty gradient is almost ten times larger andthe excess 228 Ra content shows a steep decline.
Hence we employ
different diffusivities in the two depth ranges.
Table 7-1
summarizes the diffusivities derived from this data.
As can beseen at each si te there appears to be an inverse relationshipbetween densi ty gradient and diffusi vi ty.
This is not unreasonable. The steeper the density gradient, the greater the amount ofwork against gravit~ required to mix the water!
Near-bottom 22 Rn profiles have been measured at perhaps 200
localities.
At many of these sites the density structure shows apattern similar to that seen near the air-sea interface.
Immedia tely above the sediment-water interface there is a mixed layer
ranging from 30 to 200 meters in thickness.
This mixed layer iscapped by a benthic thermocline.
In such situations the radon isusually nearly uniformly distributed throughout the benthic mixed
layer and penetrates only a very short distance into the overlying
benthic thermocline.
Thus only upper limits can be placed on the
diffusi vi ty in the benthic thermocline and only lower limi ts canbe placed on the diffusi vi ty in the benthic mixed layer.
Hence,
such regions yield only qualitative information regarding therelationship between vertical diffusivity and density gradient.
There are, however, a number of sites which show the neat
exponential gradients seen in the two examples gi ven above.
The
diffusivities obtained at these sites are plotted against thecorresponding dens i ty gradients in figure 7-5.
As can be seen,
these results fortify the conclusion drawn from the two examples.
Vertical diffusivity appears to be inversely proportional todensity gradient!
On this same diagram are given vertical diffusivities determined in ~he ocean I s thermocline using the penetration of thetritium produced by nuclear testing and those determined in smalllakes through artificial tri tium injections.
These resul ts fallalong the trend established from the 222 Rn and 228 Ra results fromthe deep sea.
The near constancy of the product of vertical diffusi vi ty and densi ty gradient sugges ts that the energy expended
against gravity in mixing less dense with more dense water has thesame order of magnitude in all these situations!
While the plot in figure 7-5 provides compelling evidence
that a major factor influencing the rate of vertical mixing in theocean (and 'also in lakes) is the density gradient, this conclusion
cannot go unchallenged.
The problem lies in one of the two
simplifying assumptions made in order to analyze the 222Rn and
228 Ra data.
Let us examine them.
The assumption regarding theneglect of advection poses no great problem.
If the r'ate of
upward advection is anywher'e near the value of about 4 m/yr
obtained from the 14 C. based ventilation time for the deep sea,
then the movement on the time scale of the half-life of radon(0.01 year's) is trivial (0.01 years x 400 cm/yr is only 4 cm).
In
the case of 22 8 Ra whe r'e the half-life is 5.5 years again theamount of movement (i.e., about 22 meters) is small compared tothe observed upward penetr'ation of 228 Ra into the water column.

327

Table 1-1. Relationship between the densi ty gradient (!J.p/!J.z) andapparent vertical diffusivity (K) at GEOSECB station 32 in thenorthwestern Atlantic and at GEOSECS station 263 in the southwestern Pacific.
As the near bottom density gradient is too small
to measure at these stations, the density gradient is computedfrom the near bottom temperature gradient (!J.pj!J.z) and the density
potential temperature relationship (!J.p/!J.8) for the deep part of
the water column.
The product of the mixing rate and the density
gradient is a measure of the work done against gravity. The 222Rn
measurements were made as part of the GEOSECS program (81).
The
228Ra measurements were made under the direction of Herb Feely at
the Lamont-Doherty Geological Observatory and of Willard Moore at
the University of South Carolina (91).

Depth Range
km

Tracer

!J.8j!J.z!J.p/ !J.8

!J.p/!J.z

K

%%C

%o/km

cm 2/sec

K (!J.p/ !J.z)°C/km

Station 32

4.1 - 5.0

228 Ra-0.35-0.074

0.026

16

0.4

5.4 - 5.9

222 Rn-0.019-0.14

0.0027400

1.1

Station 263

4.2 - 4.7

228 Ra

-0.47

-0.22

0.103

5

0.5

5.4 - 5.7

222 Rn-0.013-0.14

0.0018

115

0.2

328

•

.-..
•• "

•

222 Rn DE EP SEA~ 228Ra DEEP SEA•

3 H N.SEA THERM .

S

,~.

10

• ~+

't
BUOYANCY FLUX
4 x 10- 6 cm 2 /5 3

LAKE224~
HYPOLIMNIONLAKE 227~
HYPOLIMNIONLAKE 224
THERMOCLI NE
LAKE 227~
THERMOCLINE _10-5'--_-L._ _....--_......_

102.......~
4...______
6

10
10
10
DENSITY GRADIENT (I0-lIg/cm4j

Figure 1-5. Apparent coefficient o~ vertical diffusivity versus
density gradient based on fourteen
22 Rn profiles in the deep sea(91), for two 228 Ra profiles in the deep sea (91), and forpenetration of bomb-produced tritium into two regions of thethermocline of the Norwegian Sea (62).
Also given are the resultsfrom tritium injections made by Paul Quay of the Lamont-Doherty
Geological Observatory in two small lakes of the Experimental
Lakes Area near Kenora, Canada (67).
The solid line represents a
constant product of diffusivity and density gradient.
This
product, called buoyancy flux, is the rate at which energy isexpended against gravity in accomplishing the mixing.

329

45.89...-----.......- - - - - - . . - - - - - , . . - - -..

•
•

45.91

w
o

~ 30°

,

~

.J

600 400
LONG I TUDE ow

~

o
~
!L,

I 0~00

45.93

••

• •

45.95

*
o

•
•
••

0.5
1.0
228 Ra (dpm/IOOkg)

1.5

Figure 7-6'. Plots of 228 Ra activity versus potential density(referenced to a depth of four kilometers)
at several stations in
the western basin of the North Atlantic (96).
The station
locations are shown on the inset map.
The measurements were madeat the Lamont-Doherty Geological Observatory on samples collectedduring the GEOSECS program.

330

The trouble stems ra ther from the ne~lect of horizontal transport. It is possible that much of the
22 Rn and 228 Ra found above
the benthic mixed layer was carried along isopycnal horizons fromadjacent benthic hills (or in the case of the longer Ii ved 22 8Rafrom ve ry dis tant continental margins and ridge flanks).
One
piece of evidence in support of this hypothesis is shown in figure
7-6.
228 Ra profiles from several widely separated places in thenorthwestern Atlantic Ocean all show roughly the same concentration on a gi ven isopycnal surface.
This is the pattern to be
expected if 22.8 Ra were spread rapidly across the isopycnals inthis deep sea basin.
There is also a theoretical problem wi th the inverse correlation between diffusivity and density gradient.
If thisrelationship applies throughout the deep sea then there is noplace for upwelling.
The observed down water column decreases in
the gradients of temperature and salinity would be a consequenceof the changing density gradient.
This proves unsatisfactory for
two reasons.
First, the water descending from the deep water
source regions has somehow to return to the surface.
Upwelling is
the only way to dispose of it.
Second, if diffusi vi ty is determined by the densi ty gradient and the densi ty gradient by diffusivity then the ocean is in a rather precarious dynamic state.
The hypothesis that the profiles of temperature, salinity, anddensity are produced by the diffusion down into an upwelling plumeis far more acceptable from a theoretical standpoint.
One might then ask why horizontal processes generate such aneat correlation between density gradient and the apparent diffusivities calculated from the vertical profiles.
No one knows theanswer to this question.
We have reached the forefront of research in this area.

THE DISTRIBUTION OF HELIUM-3 IN THE DEEP PAC~IC
The third tracer which offers information with regard to the
diffusive processes at work in the deep sea is 3He.
As mentioned
above, the distribution of this isotope has something to say about
the relative importance of isopycnal and diapycnal mixing.
This
plume of 3He is well documented in the Pacific Ocean.
The ridge
crests in the Pacific are shown in the map in figure 7-7. Hot water issuing forth from these crests are thought to be the major
source of excess 3He •An east-west section of excess 3He acrossthe East Pacific Rise at 15°S is shown in figure 7-8 (the positionof this section is marked A-B on the map in figure 7-7). One lookat this diagram leaves no question wi th regard to the source ofthe excess helium.
A plume extending across the entire Pacific to
the west of the Rise is seen.
A north-south section (marked C'-D
on the map in figure 7-7), somewhat west of the crest of the East
Pacific Rise, is shown in figure 7-9.
A second north-south section (marked E-F on the map in figure 7-7), along the western partof the Pacific basin, is also shown in this figure. Although lessdistinct in the far north, the mid-depth maximum of 3He is foundthroughout the Pacific.
These diagrams make it remarkably clearthat the 3He released from the ridge crests is able to spread theentire width of the Pacific Ocean before it is carried to the

331

E

,,
\

... )
I

A --+-,.1-_ _
I

,-'
'-~

,I

I

I
I
I

1

D

,

,1' . . . . . -,-'/'
'
60·

~/

Figure 7-7. Map showing the locations of the 3 He sections shown infigures 7-8 and 7-9. Also shown are the active ridge crests.

EAST PACIFIC RISE
IOOOkmIOOOkmI

5 L--L_..I..---L_-"---L_...L-........_.....I---IL...-........- - i
130

120

110

100

LONG I TUDE ow:Figure 1-8. Section of excess 3 He content (expressed as o3 He inpercent) across the East Pacific Rise at 15° south (line A-B in
figure 7-7).
The westward extending plume of 3 He originating at
the cres t of this spreading center is qui te apparent.
The measurements on which this diagram is based were made by John Lupton
in the Laboratory of Harmon Craig at the Scripps Institution of
Oceanography (280).

332

Or--ro--.....-ro-~......

--...----I..-.......

- - - I......

E

=

2

.-a.. 3
:I:

w

o

4

5
40

......

..a..-.iI~~_ ......_

200......._

20......_ . . I

40

5
LATITUDE
N
O...-.....--...--.---....--P"'--P"'--P"'-.....- ...

-

E

-

~2

::r:

f-

a.. 3
w
o
4

5'"'-.....

_-a._-a_-~_.a.-_

40
S

20

0

. . ._.20. . _.. . -----40

LATITUDE

N

Figure 7-9. North-south sections of excess 3 H in the western (line
E-F in figure 7-7) and eastern (line C.~D in figure 7-7) Pacific.
The 3 He maximum generated by the release of hot waters from theridge crests is displayed prominently across the entire ocean.
Only in the northwestern Pacific does it lose its identity.
The
dashed lines show the depth of the ridge cres ts at 15° Sand 13° N
where intense hydrothermal activity has been observed.
The measurements on which these diagrams are based were made by Brian
Clarke at McMaster Unive~sity in Hamilton, Canada on samples collected as part of the GEOSECS program.

333

surface by upwelling.
In fact, 3He released from the ridge crestsof the Pacific reaches the deep circumpolar Antarctic and eveninto the deep Atlantic Ocean!
Thus, the distribution of excess 3He in the deep sea reinforces the suspicion that 228 Ra may move considerable distances(1000 kilometers or more) in the horizontal duri~ its nine-year
mean lifetime.
If so, the vertical profile of 22 Ra, measured at
any gi ven pIa ce on the sea floor, is c rea ted by a c omp I ex in t e raction between horizontal and vertical motions and is strongly
influenced by the topography of the basin and by the distributionof sediment type wi th depth in the basin.
The importance of the
shape of the topography can best be seen by considering how 228Ra
would be distributed if horizontal mixing was exceedingly rapid
and vertical mixing did not occur.
In this case, 228 Ra wou~<; bedistributed uniformly across each isopycnal.
The amount of
8Ra
at any given depth would then depend on the ratio of the volume ofwater between any pair of isopycnal surfaces to the area of sediment intersected by these surfaces.
This ratio would in turndepend on the shape of the basin (the ratio of water volume tosediment area would always increase wi th distance above the sea
floor but the pattern of this increase would depend on the detailsof the topography of the basin).
The mean concentration of 228 Ra
between a pair of isopycnal horizons would also depend on thesediment type intersected by these isopycnals.
As the amount of
228Ra released from the sediment would be proportional to its
232Th content, sediments rich in calcite and opal (phases free of
232 Th ) should show lower 228 Ra fluxes than red clays.
Thus theincrease in 228Ra concentration expected, due to the rise in sediment area to volume ra tio wi th depth, would be to some extent
compensated by the decrease in calci te content of the sedimentwith depth.

SOURCES OF DEEP WATER
Deep waters are being formed in two regions of the ocean: inthe northern Atlantic and along the perimeter of the Antarctic
continent.
In the discussion which follows we will designate theformer as the northern component and the latter as the southerncomponent to global deep water.
The proportions of these twocomponents vary from place to place in the deep sea.
As the twocomponents differ in their concentrations of the tracers ofinterest, part of the geographical variation in the properties ofabyssal waters is the result of the differences in the proportionsof the two components.
If we are to make effecti ve use of thesetracers, we must develop a means of determining the relative
contributions of these two water types to a given water sample.
Thus, we must learn more about the conservative* properties ofthese two components.
The establishment of the compositions of these two watertypes is more difficul t than one might suspect.
The reason is*A conservative property is one that does not change during residence in the deep sea (i.e., one that is not altered by respiration, solution, or radiodecay).

334

that the process of deep water formation is quite complicated.
The cooling which triggers the descent of new water into the abyss
occurs during very brief periods at winter's peak. Oceanographers
find work in these regions at this time of year very difficul tbecause of darkness, cold, ice and frequent storms.
Furthermore,
the actual sinking process is confined to small areas whose locations are not well known and may vary from year to year.
Thus,
the chances of finding and properly documenting them is difficult(like sitting in a very large bathtub with ten drains waiting todescribe the action when, by surprise, one of the plugs ispulled) .
Even if adequate studies could be made of these plumes 0 f
descending winter water, we still would not have our answer.
As
the waters descend they entrain water through which they pass.
In
some cases the water first reaches a temporary holding basin.
From here it cascades over a barrier into the open ocean.
Whilein the holding basin it may lose oxygen, gain nutrient constituents...
As it cascades from the holding basin into the open sea
more entrainment occurs.
Thus, our choice of an endmember composition becomes somewhat a matter of definition.
We must sample
water far enough away from the source regions that the cascade
entrainment has already occurred, but not so far away that mixing
with the "old" waters present in the basin has introduced waterfrom the other source.

NORTHERN COMPONENT WATER
Northern component water is a mixture of water from three
distinct sources. The pathways followed by these waters are shownin figure 7-10. Two of these waters have as their ultimate source
waters which sink in the northern part of the Norwegian Sea.
This
water is held behind the sills which connect Greenland to Icelandand Iceland to the Bri tish Isles.
The bottom water behind thesills has a temperature of about -1.0 o C: and a salini ty of about
34.91 0 /00.
Periodically, (presumably mainly during the winterwhen more new water is descending into the Norwegian Sea holding
basin), this deep water spills out over the connecting sills andcascades into the deep Atlantic.
That water spilling over thesill to the west of Iceland enters directly into the head of thewestern basin.
This water, which is called the Denmark Straits
Overflow Water (DSW), has the highest density of the three contributors to northern component water (see table 7-2). Water also
spills over the sill to the east of Iceland.
This water entersthe eastern basin of the Atlantic.
CDriolis forces (i.e., forces
resulting from the rotation of the earth) cause this descending
cascade to hug the margin of the Mid-Atlantic Ridge and then toflow through the Gibbs Fracture Zone (a deep gap in the Mid-Atlantic Ridge) into the western basin.
Surprisingly, this water does
not ventilate the abyssal eastern basin of the Atlantic but passes
almost entirely into the western basin. *This second component*The eastern basin is ventilated by western basin deep water which
passes through the Ridge via another major gap called the Romanche
Fracture Zone. This gap is located just south of the equator.

335

80°

60°

40°

20°

20°

70°

• I

50°

Figure 1-10. Map showing the paths f'ollowed by the deep waters
t'ormed in the northern Atlantic.
Water from the surface of' theLabrador Sea sinks to the bottom of the western arm of the westernbasin forming LSW.
Water trapped behind the ridge connecting
Iceland with Greenland spills over the sill into the eastern armof the western basin forming DSW.
Water trapped behind the ridge
ooonee tlng Iceland and the Sri tlsh Isles spills over the 5111 andflows along the east side of the Mid-Atlantic Ridge.
When itreaches the Gibbs Fracture Zone, it passes through the Ridge into
the eastern arm of the western basin.
Because it 1s less dense,
GFZW overrides the DSW flowing along the bottom of this arm of thewestern basin.
These waters mix to form northern component water
which enters the main western basin along the continental marginof North America.

336

of northern component water is called the Gibbs Fracture Zone
Water (GFZW).
It has a lower potential density than DSW (seetable 7-2).
The third contributor to northern component water iswater which sinks from the surface of the Labrador Sea during thewinter.
This water descends directly into an arm of the westernbas in rather than via a holding basin as do the other two.
Itsdensity is less than those of the other two (see table 7-2).
One might ask how we know that these are the chief and only
contribu tors to northern component waters.
Ou r bes t evidence
comes from plots of e and S for deep waters from the far northern
Atlantic.
As shown in figure 7-11, these wa ters occupy a triangular region on such a diagram.
The composi tions of the waterat the three corners of this triangle correspond to that of water
filling the Gibbs Fracture Zone, to waters flowing along the bottom south of the Denmark Straits and to deep water in the LabradorSea.
The waters within the triangle are presumably formed bymixing these three endmembers in various proportions.
A point
near the center of the triangle would correspond to a mixture
consisting of roughly equal proportions of the three endmembers.
A point half way along -one of the s ides of the triangle would
correspond to 50 - 50 mixtures of two water types defining thisedge.
While it is possible that other water types with compositions lying within the triangle also contribute, this is unlikely
since source regions for such waters have not been identified.
A traverse of stations down the northernmost part of thewestern basin (see figure 7-12) reveals that the mixing along
density surfaces is qUickly accomplished (rather than having waters with a variety of temperature-salinity combinations only anarrow range of temperatures are observed on any given isopycnal).
Further, the density of the bottom-most water decreases tothe south.
This presumably is the result of vertical mixing.
As
this traverse is extended further south, mixing with "old" waters
in the deep-western basin becomes significant.
This is best seen
on a plot of dissolved silica versus salinity (see figure 7-13).
While all three contributors to northern component water have low
silicate contents, waters to the south of SooN have progressively
high silicate contents.
This increase is primarily the result of
the intrusion of southern component water~
To understand this argument, we must look more carefully at
the distribution of silicate in the Atlantic Ocean.
The silica
content of the three of the contributors to northern component
water are as follows: 11 )Jm/kg for' LSW, 11 )Jm/kg for DSW and 14)Jm/kg for GFZW.
By contrast, the silica content of Antarctic
Bottom Water (AABW) entering the abyssal western Atlantic from thesouth is 125 )Jm/kg.
Because of mixing with AABW (and perhaps also
because of opal solution), the deep waters in the central Atlantic
contain significantly higher amounts of silicate than northerncomponent water. Because of this, increases in silica content area good tip-off that mixing between "new" and "old" water has
occurred.
If we assume that the sole source of the excess silica is
mixing with AABW, the temperatures and salinities for the northern337

....---....-- .....- -.--•. .. ••••••••••• .. - • -• . .., .. .' .. .
•........,

W

::>

•

• ,e

• .... • •

:.~

~

'

y

., •

"~D~l~~

45JO

a::
w 3
a..

• ••
• •••

•••
A:"
~ _•• •.
,..
'..: :t -,,!I '.;.. .', : "-" • •.: J.. :.\1&.. l\~ ~'1.,1., ..... : ·. \ .

: •••• : .'fe. '. • •

0::

.-«

• ••

•

u
o

45.75

~

45. 80

I-I

45.85

w

«

I- 2
Z
W

.-o

a..
-34.8535.00

Figure 7-11. Diagram showing the potential temperature-salinity
relationship for waters from the northernmost part of the western
Atlantic (i.e., > 50 0 N).
Contours of potential density referenced
to 4000 meters are also shown.
The circle at the right designates
the composition of water found in the Gibbs Fracture Zone; theupper circle designates the composition of deep water in the
Labrador Sea and the lower circle waters generated by thespillover at the Denmark Strai ts.
Points within the triangle
defined by these three circles are mixtures of these three end
members.
Based on a cornpila tion of the da ta for many expedi tionsin the Atlantic Ocean made by Reid and his colleagues at the
Scripps Institution of Oceanography.

338

4,.....--.,...--......

....-__......-.

LSW

w
a::

::)

~
w

3

a::

C-

~

.-

/

W

-l

<t: T0GFZWNORTHERNCOMPONENTWATER

2I- AABW

o

Z

W

~

I-

~

•
•
•

DSW

:3 (SION)
2 (48°N)
I (45°N)
27 (42°N)
28 (:39°N)

I

34.90

34.95

35.00

SAL IN ITY (%0)
Figure 7-12. Potential tempera ture and salini ty of the northerncomponent wa ters found in the northern part of the Atlantic's
western bas~n (i.e., 39°N to 51 0 N). A correction has been applied
to remove the influence of the small amounts of AABW wa ter foundat these latitudes.
The potential temperature of AABW is -O.loC
and its salinity is 34.67 0 /00.
This correction is based on theexcess silicate concentration in the water (southern component
water carries 125 ~m/kg of silica and the northern component
wa ters only 12 ~/kg).
The mos t dense northern component wa terreaching the main basin of the central North Atlantic has apotential temperature of about 2.1 o C and a salinity of about
34.93%0.
The arrow extends from the composition of this water
toward the coordinates of AABW.
The points shown are for GEOSECS
stations (424).

339

Table 7-2. The compos i tions of the wa te r'S trapped to the north of the topographic
barrier connecting Greenland with Iceland, and Iceland with the British Isles, thecompositions of .the three contributors to northern component water,
and thecomposition of the most dense norther'n component water reaching the western basin of
the North Atlantic.
Based on measurements made as part or the GEOSECS program(21, 424). Recent discussion of the origins of these waters can be found in references 724, 725 and 726.

e

S

04*

NO 3

P04

SiO 2

AOU

~ 14 C

°C

0/00

0/00

~m/kg

~m/kg

PO 4°

~m/kg

~m/kg

~/kg

0/00

-1.04

34.913

46. 46

16.1

1.02

14

53

0.62

-67

Sea Bottom-0.93

34.920

46.45

14.9

0.94

11

59

0.56

-

Denmark Strai ts
Overflow

1.79

34.915

45. 99

15.1

0.97

11

43

0.65

-

Gibbs Fracture
Zone

2.81

35.000

45.86

16.2

1.07

14

50

0.71

-

Deep Labrador Sea 3.50

34.945

45.68

17 .6

1.10

11

43

0.79

-

Densest NCW
entering Western Atlantic

34.932

45.93

16.1

1.04

12

50

0.73

-67

Wa ter TypeE. Norwegian
Sea Bottom(..A,,),+:a..

ow. Norwegian

2.10*04 is the potential density referenced to a depth of four kilometers.

component water can be computed.*
In figure 7-12, the temperatures and salinities for GEOSECB stations 1, 27 and 28 have beencorrected in this way for the contribution of AABW.
We can see
from this diagram that the highest density northern component
water reaching the north central Atlantic has a potential temperature of about 2.1°C and a salinity of about 34.93 0 /00.
These
corrected e-S values fall along the trend seen at GEOSECS station3, where little excess silicate is present (see figure 7-13).

C.DNSERVATIVE PROPERTIES OF NC~
If we are to trace the contribution of northern component
water to bottom waters throughout the world ocean we need sometags which will distinguish it from southern component water andfrom water mixed down from intermediate depths into the deep sea.
Salinity is the most reliable tag in that no process occurring in
the deep sea can alter it significantly. Temperature is not quite
so conservative a property. Geothermal heat from the interior of
the planet i~ entering the sea floor at the rate of about 50 calories per em per year. In 1000 years this heat flux would raise
the temperature of a water column two kilometers in height by
O. 4°C!
Hence, mixtures of bottom water at various places on thesea floor will be slightly warmer than expected from the end
member temperatures.
The range of potential temperature (2°C) and of salinity(0.OB5 0 /00) encompassed by the three contributors to northerncomponent water is quite large.
While in the discussions which
follow we will assign this water the e and S found for the densest
mixture of these contributions reaching the open Atlantic (i. e. ,2.l o C and 34.93 0 /00), it must be remembered that northern component water has the range of e and S values shown in figure 7-12.
*These_ temperatures and salinities are calculated from the following equat ions:
125-[H4SiO It]
125-12S-(l-f n )34.667
fnwhere f n , en and Sn are, respectively, the fraction, potential temperature and salinity of the northern component contribution to the mixture, e, Sand [H4Si04] are the measured potential
temperature, salinity and silicate concentration in the mixture.
The characteristics of AABW are e = -a.loe, S = 34.667%0, and[H4 Si0 4] = 125 ]..1m/kg. The mean [H4Si04] content of northern component water is 12 ]..1m/kg.

341

40

TO
AABW

-

30

Ql

~

••
".

""'0

:l
.....,
q-

0

.-

en

..
..... e.

...:.

E

20

v

• 121

:r:

•
•

28
27

0
0

3
4
5

•
* 2
I

6

10

0

*

.

200

...•
~

.0

•

10

t~~i· @
~p~

GFZW

8
24

34.85

.

34.90

34.95

35.00

SALINITY (%0)
Figure 7-13. Plot of silica versus salinity for deep waters in thenorthern Atlantic.
Wa ters to the north of 50 0 N have silica tecontents less than 15 ~m/kg.
To the south a rapid rise insilicate content is noted.
The samples with high silicate content
have 8-8 coordinates falling outside the triangle of northerncomponent compos i tions.
These devia tions can all be explained
through the admixture of AABW.
The arrow extends from thecomposition of the most dense northern component water towards thecomposition of the bottom water entering the South Atlantic
through the Vema Channel (8=34.67%0 and H 4 Si0 4 =125 wn/kg).
The points shown are for GEOSECS stations (424).

342

Another nearly conservative property is what chemical oceanographers term preformed phosphate (i.e., the amount of phosphate
carried by the water when it leaves the sea surface). This property is computed as follows:
AOU[P0 4 J O = [PO]4 in situ - 135

7-5

The amount of O2 consumed in the water (i.e., its AOU) is used tocorrect for the contribution of respiration to the phosphate content of the water.
As we have seen, while the ratio of P04 generated to oxygen consumed averages 135, it does not always havethis value.
Hence PO 4° calculated, using a fixed ratio of 135
will not be exactly conservative.
The higher the AOU value for agiven water sample, the larger the uncertainty in P04°.
As
shown by the P04° versus depth plots in figure 7-14, this property
does have considerable potential.
Water formed in the northern
Atlantic has a considerably lower P0 4 ° content than does waterfrom the Antarctic Ocean! As shown in table 7-2 the range in P04°
values for the three contributors to northern component water isquite small.
As we shall see later in this chapter the H/D and 18 0 /16 0
ratios in the water itself are excellent conservative tracers.
However, as not enough analyses are available, only preliminary
use can be made of the distribution of this property in global
bottom water studies. Hence, they are not included in table 7-2.

INITIAL RADIOCARBON TO CARBON RATIO IN NeW
If we are to use the distribution of radiocarbon to estimate
the ventilation time of the deep sea, we must establish the initial 14 C/ C ratio for northern component water.
This presents a
problem because at the time of the GEOSECS survey virtually all ofwhat we term here as new northern component water had been contaminated with 14 C produced during nuclear testing. Fortunately, it
is possible through the simultaneous consideration of tritium
measurements (the tritium present in the ocean is almost entirely
from nuclear testing) and radiocarbon measurements to sort out
this contamination.
As shown in figure 7-15 waters from the Norwegian Sea and northern Atlantic (>50 0 N) show a fairly tight correlation between their tri tium and radiocarbon contents.
Thosewaters in this region which had not as yet been contaminated with
bomb-produced tritium (the prenuclear testing tritium content ofwaters in thi y region was likely on the order of a few tenths of aT. U.) have b. 4 C: values averaging about -670/00.
This value istaken to be that carried by the carbon in these watEfrs prior tothe onset of nuclear testing.
Al though we gi ve a !:J. 4 C: value intable 7-2 only for mean northern component water, it is likelythat all three contributors had nearly the same value.
FEED FOR NC}/ PRODUC?rIONAs

can beseen from

table

343

7-2,

the

three

contributors

toPO'; (~mol/kg)
1.0

0.5

0

-E

.:I:

0

1.0

0.5

1.5

0

2

0
0
0
0
0
0
0
0

0

0.5

SOUTH
ATLANTIC
18°S-31°W

0
0

1.0

0

0

0.5

1.0

'000

1.5

0

0

0.5

0

0
NORTH
PACIFIC
31°N-17JoE

00

7-14. Plots

~

0

0
0
0
0
0
0

0

0

DRAKE
PASSAGE
58°S-66°W

0
0
0

0
0

5

1.5
0-'0 c§J
<:)
00

,

0

3

1.0

0

\

2

0

1

8

E

ocean.

0

P04 (,u.mol/kg)

0

5

0

0

~

0

NORTH
I NOlAN
6°N-64°E

0

<:)

5

P04 (,u.mol/kg)
1.5

0.
0
0

SOUTH
ATLANTIC
45°S-51°W

50

AOU/I35)

0

0
0

B

0

Figure

.

80

.~

4

0

~

P04 (,u.mol/kg)
00

1.5

.

2

0
0

5

W

09
1.0

0
0

NORTH
0
ATLANTIC ~3SON-47°W 0....Ia.

0.5

0
0
0

W

.:I:

0

0

0

....Cl.J:-

P04 (,u.mol/kg)

P04 (fLmol/kg)

0
0
0

~

5

of ini tial PO 4 content (i. e., PO 4measured versus depth for key localities throughout the world
Based on da ta obtained as part of the GEOSECS program(424,425,426) •344

75..--.....-....--.....- ......--,~-'r---.....--..,

50---...
0

-v<J

•

• fI'

25

0~
........,

U...I.-GEOSECS RESULTS>50 o NATLANTIC OCEAN

0

•

-25

.
~­
- ..I-.-~••••<:) DEEP NORWEGIANSEA• OTHER2

6

8

Figure 7-15. ~14C versus tritium content for waters from theno rthern A~lantic in 1972.
The trend termina tes in wa ter wi th atritium content less than 0.2 T.U. and ~14C values close to-67%0.
The tritium resul ts were obtained by Ostlund of the
University of Miami as part of the GEOSECS program (51).
Theradiocarbon resul ts were obtained in the laboratories of Ostlund
and Stuiver (21).

345northern component water have moderate nutrient constituent contents and significant oxygen depletions.
Two conclusions can be
drawn.
First, the waters feeding deep water production in thenorthern Atlantic come on the average from a depth of at least 700
meters in the temperate Atlantic.
Were this not true, the outgoing new deep water would be carrying more silicate, more phosphate, and more ni trate than was carried by the incoming feedwater.
This would require a nutrient source within the Norwegian
Sea far larger than that constituted by all the world's rivers put
together!
As there is no possible source of these consti tuents
within the Norwegian Sea, we can say with confidence that the concentration of these substances must be equal to or possibly
slightly greater* in the waters feeding the deep water formation
process than in the new deep water itself.
To meet this requirement the water must come mainly from depths greater than 700
meters (see figure 7-16).
Second, the fact that this new water begins its trip throughthe deep sea wi th a sizeable AOU suggests that plant material
formed in the water during its residence at the surface in the
Norwegian and Labrador Seas is destroyed by respiration in thenorthern holding basins.
As the preformed phosphate content of
the eXiting water averages 0.7 ~m/kg, while the measured phosphate
content is 1.1 ~m/kg, about one third of the phosphate must havebeen utilized by plants as the feed water underwent cooling.
This
plant matter then falls to the holding basin where it is largely
eaten by animals and bacteria.
The most likely candidate for the feed water is the the silicate maximum water.
As shown in figure 7-16 this water> is foundat about one kilometer> depth.
This featur>e is likely generated by
the isopycnal transport of silicate r>ich water all the way fromthe Antar>ctic Ocean!
In fact the reason for the existence of thismaximum is likely the withdrawal of water from this density levelfor conversion to deep water.
This withdrawal permits the northward penetration of Antarctic der>ived waters (at other depths in
the North Atlantic, water is being exported to the south, drivingback the intr>uding Antarctic waters).
The fact that the feed water for the source regions of northern component water is from intermediate depth rather than fromthe surface accounts for its rather low ~14C value, (surface waterin the temperate zone had a ~ 14 C of abou t -45 0 /00 prior to theyear 1950),while new northern component water had a ~14C of about-67 0 /00.
As shown in table 7-3, based on the few 14 C measurements made· on water from intermedia te depths made prior to the
input of 14C from bomb testing, and from the results of the
GEOSECS expedition cor>rected (using tritium data) for the contribution of bomb- 14 C, a 6 14 C value of about -890/00 is obtained
for the feed water.
The difference between the -890/00 value for the feed
*It is possible that some of the nutrient constituents carried
into this region are lost to the sediments, as the result of plant
growth in northern surface waters.
However, as the recycling of
nutrient elements in the sea is very efficient, this loss must be
sma 11.

346

12
14
OA.....--....--r--...---i-----......--.----.r----t

P04
NO.

0.5

COMP.

J:
~

a..

w

Q

1.0
30 0 N 39°W
CENTRAL NORTHATLANTIC........

1.5 ~-....._ ....._-~- .....--.l----~.0
.2
.4
.6
.8
1.0
1.2
1.4

P0 4 (fLmo 1/ kg)
Figure 7-16. Plots of P04 and H4 Si0 4 content versus depth in theupper waters of the central North Atlantic.
The arrows designate
the concentrations of these constituents in northern component
water.
The horizontal line segments at the heads of these arrow s
designa te the depths in the north tempera te Atlantic at whichthese concentra tions are found.
These resul ts were obtained aspart of the GEOSECS prog ram (424).

347

Table 7-3. Radiocarbon results on water samples from intermediate
depths in the temperate North Atlantic.
The first four samples
were collected before 14 C from nuclear tests had measurably
contaminated the ocean.
The last five collected during the
GEOSECS program had small amounts of tritium, necessitating acorrection for bomb- 14 C. The measurements on the samples from the1950' swere madebyBroeckerofLamont-Doherty
GeologicalObservatory (4); those on samples from the 1970' s were made by
Ostlund and Stuiver as part of the GEOSECS program (21).

Collection
Date
10/56
4/56
11/57
12/55

Latitude

oN

24
22
20
19

Longitude

oW

Depth
m

24
54
49

1100
820
950
1100

80

3

H

T. U.

-

-

-

14

b
C
0/00

-107
- 95
- 79
- 91
Mean - 93

3/73
3/73
3/7'3
9/72
9/72

33
31
28
21
27

57
39
26
54
54

1100
1080
900
1000
1010

1.05
0.30
0.15
0.10
0.25

- 82*
- 74*
- 83*
-100*
- 86*
Mean - 85*Corrected for contribution of bomb-produced 14C assuming a14
20 0 /00 change in b C per 1 T.O. increased in tritium.

348

water and the -67 0 /00 value for new bottom water must be theresult of the invasion of 14C from the atmosphere into the northern waters.
As we sha~l see below the flux ~f ~orthern component
water is about 13 x 10 m3 /sec (i.e. ~4 x 10 1 m /yr). The volume
of water contained in the formation region (i.e., north of 55°N)
is about 2 x 10 16 m3 •Thus, the flushing time of water in thesource region is about 50 years.
During this time 14C dec~~s byabout 6 0 /00.
Thus, in the absence of gas exchange the /1 C ofwater leaving the basin would be 60 /00 lower than that for input
water.
Actually, we observe it to be 22 0 /00 higher.
So exchange with atmospheric CO 2 mus t raise the /114 C value by abou t
28 0 /00.
The area of the Atlantic north of 55°N is about 8 x
10 12 m2 .
If we take the CO 2 exchange rate to be 20 moles/m 2yr andthe LCO 2 content of these raters to be 2.1 moles/m 3 then gas exchange should reduce the /1 4C difference between water and air byabou t 18 0 /00.
While this is not qui te enough the resul t lies
within the range permitted by the uncertainties that exist in many
aspects of this argument. The calculation does make it clear thatthe water does not remain in the source region anywhere near longenough to achieve equilibrium with the 14C/12C ratio in theatmosphere. *SOUTHERN COMPONENT WATER
The second major source area for bottom water in the ocean is
the Weddell Sea (see figure 7-17 for location).
DJ.ring wintermonths, water which has upwelled from the deep circumpolar current
to the surface in the Weddell Sea is cooled by contact wi th the
air and made more saline by the formation of sea ice.
As these
processes are the most intense on the continental shelf (becauseof its isolation from mixing with deep water) the most densewaters are produced along the southern and western margin of thebasin.
Under the influence of gravity, these waters cascade overthe shelf edge and into the less dense deep water.
Along theirtraverse they entrain a mixture of winter surface water from theinterior of the Weddell gyre and underlying deep circumpolarwater.
Diagrams showing plots of potential temperature versus
salinity, dissolved oxygen versus salinity, hydrogen isotope ratio
versus salinity and oxygen isotope ratio versus salinity for
Weddell Sea water is shown in figure 7-18.
In all cases Weddell
Sea water can be seen to be a mixture of about 60% surface waterand 40% entrained deep water.
The surface water component
consists of.a mixture of shelf water and of gyre surface water.
The properties of these waters are summarized in table 7-4.
The Weddell Sea bottom water has a high silica content, a rather
high AOU,
a high preformed phosphate content and a low 14C/ C
ratio.
These properties are related to two characteristics offormation process.
First, the source of the water is old deep*The m1~n ~epth of the water in the source region is 2 x 10 16 m 31
8 x 10m or 2500 m. Thus, about 5250 moles of C02 are contained unde~ the average square meter of sea surface. At a rate of 20
moleslm yr over a period of 50 years only about one carbon atom
in five is exchanged!

349

130

120

90

iI
40

\

\
I

I
I

140

II
I

150

160

10

170

- - - - - - + - - - - 1 180

10

170

20

160
, ..

30

~

",j
/

/

150

'~

"-

140

50

60

70

90

100

110

130

Figure 7-17. Map showing the pattern of deep water flow in the
Antarctic Ocean.
Northern component water joins the circumpolar
flow at the south end of the Atlantic's western basin.
Southerncomponent water is formed in a large embayment in the Antarctic
Continent known as the Weddell Sea.

350

Tab1e 7-4 CDmposition of the waters contributing to the formation of Weddell Sea Bottom
Water and of Weddell Sea Bottom Water itself.
What we designate as southern component to
deep Atlantic water is Weddell Sea Bottom Water further contaminated with ambient circumpolar water and with the products of respiration and opal solution.
Based on measurements made as part of the GEOSECS program (21,424) and by Weiss and his colleagues (428).

e

S

°C

0/ 00

Circumpolar Deep

0.62

34.708Western Shelf **-2.0

Winter Surface **-1.8

Wa ter Type(..A.)

llm/kgP04
llm/kgSi02
llm/kgAOD
llm/kgP04°
llTn/kg

46.030

32.1

2.20

123

137

1. 16

34.7

-

27.5

2.05

70

45

1.8

34.4

-

27 .5

2.05

70

45

1.8

34.64746.234

31.8

2.14

112

102

1. 28

04

N03

0/00

~ 14 C

0/00

CJ1
I--'

Weddell Sea Bottom -0. 89

-163
-146*

*Calculated rather than observed.
**Collected beneath the summer surface water.

-154

CI RCUMPOLAR....'\ DEEP

U

~

W

, WWATERa::.. (_.,

:;:)

I<t
0::

.".,

0

::E

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,:- ~ESTERN~. • ~"....·WATER'" SHELF.' r;.,. ·':~1WSBW>0

WESTERN~....._C)X·l. ,.
I- ·

lJ.J
I-

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z

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S(~WATER\ ,OEEP\', I WATER..... ...,....1 ' /WINTER
WATER

34.5

34.5

35.0

SA LI N ITY (°/00)

35.0

SALJNITY (°/00)

0.0

0
CIRCU
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1•-

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r~lWS8W
\/1

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/~I
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i~ I
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35.0

SAL IN ITY (°/00)

34.5

35.0

SALJ NITY (%0)

Figure 7-18. Plots of potential temperature, dissolved oxygen,
18 0 / 16 0 ratio and D/H ratio agains t salini ty for wa ters from theWeddell Sea.
In each diagram the composition of Heddell Sea
Bottom Water (WSBW) falls within the triangle defined by thecomposi tions of winter surface wa ter in the open Weddell Sea,
western shelf water from the margin of the Weddell Sea andcircumpolar deep water entering the Atlantic sector of the
Antarctic Ocean through the Drake Passage.
These resul ts wereobtained by Ray Weiss and Harmon Craig at Scripps Institution of
Oceanography, La Jolla, California (428)_

352

wate~ which is being ~ecirculated around the Anta~ctic.
Second,the residence time of this water at the surface is so short thatplants cannot fully utilize the nutrient elements al)d that gas
exchange can make only a small dent in restoring the
4 C/ C ~atioto that for atmospheric CO 2 •The salinity of the water changes
only a small amount (from precipitation and continental ice melting).
The main thing which happens is that the water is cooled byabout 1.5°C as the result of contact with the atmosphere.
The newly formed Weddell Sea wate~ joins the circum-Antarctic
deep flow and mixes upward into the older water producing thelinear arrays of 8-S seen in figure 7-19, of 8-D/H seen in figure
7-20 and of 8-PO 4 ° seen in figure 7-21.
As these distributions
all terminate at the composition of Weddell Sea Bottom Water(WSBW); other sources of deep water around the perimeter of the
Antarctic are either too weak to make an impact on these trends orthey have composi tions falling along the main trend.
The 6 14 C
values for these main sequence waters are remarkably uniform (seefigure 7-22) .
This is not unexpected in the light of the factthat Ii ttle 14 C renewal occurs during the cooling process in the
Weddell Sea.
It is this main sequence of Antarctic waters which feeds
abyssal waters into each of the major oceans.
In each case entry
is via a barrier which partially blocks communication between the
Antarctic and the ocean basins to the north.
In the western Atlantic it is the Rio Grande Rise located at 31 ° 8 which blocks
communication. A deep cut in this barrier, the Vema Passage (VP),
permits bottom water to spill through to the north.
In the western Indian Ocean the Southwest Indian Ridge , located at 32° S,
partially isolates Antarctic waters from those in the westernbasin of this ocean.
The Atlantis Passage CAP) provides a gap
through which deep water can move.
In the western Pacific the
Samoan Ridge located at
10° S provides the impedance to mixing.
The Samoan Passage (SP) provides the conduit for the deep water
entering the equatorial and northern Pacific. The compositions of
the waters spilling through these three passages are shown infigures 7-19, 7-20, 7-21 and 7-22.
As can be seen they all lie
along the main sequence of circumpolar deep water compositions.
Once beyond the passages these bottom wate~s begin to mix
with the less dense overlying water in the major ocean basins.
In
the Atlantic and in the Pacific these waters have composi tions
which fall well off the main sequence for circumpolar waters.
Thus, in these two oceans we see qui te a change in the 8-8 andP0 4 o -S relationships from those observed in the Antarctic (seefigures 7-23 and 7-24). For the Indian Ocean these relationships
are more similar to those in the Antarctic (bearing witness to theabsence of a deep water source at the north end of this ocean).
As shown in figure 7-25, the radiocarbon-salini ty trends in
the Pacific and Indian Oceans are qui te different from those in
the Atlantic.
In the Indian and Pacific Oceans radioacti ve decay
dominates, leading to an up-wa ter-column decrease in 1114C~.
By
contrast, in the Atlantic mixing between the Vema Passage Waterand the high 11 14 C northern component water generates and upwater-column increase in 11 14 C.

353

2,---,,.--,.--,,.-.....,r--,---,.---..---..---..---..-.....
INDIAN 0 430 60°5
SECTOR El 431 64°5
A 432 59°5

W
Il:
::J

r<!

a:

I

W
0-

PACIFIC. 282 57°5
SECTOR • 287 69as
• 290 58°5
78 61 °5
ATLANTIC ED
82 56 as89 60 aS
SECTOR ~

*

~

W

r-I

~ 0

rz
w

f-

a::
34.66

34.70

34.68

34.72

34.74

SALINITY (%0)

Figure 1-19. Potential temperature versus salinity for the "mainsequence" of circumpolar deep waters.
The coldest and least
saline waters on this sequence correspond in composition to
Weddell Sea Bottom Water.
It is this water that we designate
southern component water.
Also shown are the compositions of thewaters spilling northward out of the Antarctic into the western
Atlantic (via the Vema Passage), the Western Indian (via the
Atlantis Passage) and into the Pacific (via the Samoan Passage).
The locations of these key passages are designated on the foldout
map.
The data shown here were obtained as part of the GEOSECS
program (424,425,426).

lJ.J

a:
:::>

f-

<!

a:
lJ.J
c:E
lJ.J

f-

~
_

I-

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A

0;;:
.
0

v·
.

.

V]

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.

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PACIFIC. 286 66 0 5
293 53°5076 58°5

SECTOR •8~

ATLANTIC ~SECTOR

0.. -I .' - _......_ _....-._........

-3

-2

&
\oJ

~

-I

8D (%)

78 6/°5
67 45°5
82 56 °5

85 58°S
89 .......
60 0 S_~a

Figure 1-20. Potential temperature versus D/H ratio (given as cD)
for circumpolar deep waters.
The linear array terminates at thecomposition of Weddell Sea Bottom Water.
The isotope results wereobtained in the laboratory of Craig of Scripps Institution for
Oceanography on samples collec ted as part of the GEOSECS program.

354

*.&
<...>

~

•

W
0:

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0::

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a..

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w

18 6'°5
82 56°5

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.,.

mA~\

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ATl. EDA.- SECT.

ZWo.& 290 ~8"S

~~0!)
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<t

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~i •• tr

INO. 0 430 60"5
SECT. G 431 64"5A 432 59°5
PAC. • 282 ~7·5
SECT. • 287 69°5

II:>

:::J

J--l

0

J:) ~

m 0

~
-I L . - _ - - L_ _I - - _ - - L_ _

.....

~-

I. 2

1.0

1.4

PO; (pmol/kg)

Figure 7-21. Potential temperature versus preformed phosphate content for circumpolar deep waters.
The compositions of the waters
flowing through the Vema Passage, the Atlantis Passage, and the
Samoan Passage are also shown.
The data shown were obtained aspart of the GEOSECS program (424,425,426).
I-

PACIFIC. 282 57"5
SECTOR • 287 69°S.. 290 58"578 61°S
89 600SATLANTIC ~SECTOR

a..,
w

a:::

::>

I-

a:::
W
a..

,

I

INDIAN 0 430 60·S
SECTOR 0 431 64°S
~ 432 59"5

U

<I:

I

21-

I -

w

.

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Q..

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0,0

,

-

-150

~14C (0/00)

F1~ure7-22. Potential temperature versus 14 C/ C ratio (given asC) for "main sequence" wa ters in the circumpolar region.
The
composi tions of wa ters passing through the three major passages
connecting the Antarctic with the major oceans are also shown.
Asthe error on an individual radiocarbon measurement is ± 4%0,much of the scatter is experimental.
The radiocarbon measurements
were made in the laboratories of Ostlund of the university ofMiami and Stuiver of the University of Washington (21,22,23).
I:i 1

355

~

u
w

~

2

0::

:::>

l-

.. •

e:{

0::
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n..
~
w

c:> 010 0()/o.O

I-I-«
t-

ATLANTICA

••

Z

W

I- 0
0

37 12.0 0 N
42 1.0oN
55 18.0 0 S

INDIANEl418 6.2°N
o 425 17.3°S

a..

PACI F IC ® 226 30.6°N
E9 229 12.9°N*246 0.0°

34.6

34.7

34.8

34.9

SALINITY (%0)
Figure 7-23. Potential temperature versus salinity diagram for theworld'sdeepwatersofpotential
densitygreaterthan
45.85%0.
The stippled band represents the compositions of"main sequence" water in the Antarctic.
The trends along thewestern sides of the three major oceans are also shown.
The data
shown were obtained as part of the GEOSECS program (424,425,426).

356

1.6..-.----.~---..---...--.......- -......- -..........-----,

1.4-. ....

1.0
o~

o
a..

ATLANTICA

0.8••

37 12.0 0 N
42 '.ooN
55 18.00S, NO I AN EJ 418 6.2°N
o 425 17.3°S
PACI F IC 0226 30.6°N
ED 229 12.9°N
0
246 0.0

,•

••

••• •
••

*

0.6 L...._--L_ _- L ._ _......_ _..a-.._.-...l......_---a._ _~
34.6
34.7
34.8
34.9
SA LIN IT Y (CYoo)
Figure 7-24. Initial phosphate concentration versus salinity for
the world's
deep watersof potential density greater than
45.85 0 /00.
The
cross-hatched
area
shows
the
range
of
compositions for the Antarctic "main sequence" waters.
The
Antarctic, Atlantic and Indian Ocean waters fall along a line
connecting .the composi tions of northern component (S = 34.935,
P04° = 0.73) and southern component water (S = 34.66, PO q O1.35). The Pacific water deviates from this trend suggesting thateither a third endmember with a lower salinity and lower initial
phosphate concentration is mixing with the bottom water entering
the Pacific through the Samoan Passage or that the coefficient
relating to AOD and PO q deviates from 135 in the low oxygen deep
waters of the North Pacific Ocean.
The data shown were obtained
as part of the GEOSECS program (424,425,426).

357

•

"
."
:1·

-100

•• • ••••........
0

MAIN SEQUENCEANTARCTIC

~

0

'-"""

v

U

-<J
-200

•
0

ATLANTIC
INDIAN* PACIFIC-250 ....._ .......
34.6

34.7

34.8
SALI NI TY (%0)

34.9Figure 7-25. Plot of 14 C/ C ratio (expressed as ~14C) against
salinity for the world's deep waters with densities greater than
45.85 0 /00.
The cross-hatched area represents the main sequenceof Antarcttc waters.
The trends in the three major oceans are
generated by a combination of mixing and radioactive decay.
For
the Atlantic, mixing dominates.
For the Indian and Pacific, thedecreases below the main sequence value must be due to radiodecay.
The radiocarbon measurements were made as part of the
GEOSECS program in the laboratories of Ostlund of the Universityof Miami and Stuiver at the University of Washington (21,22,23).

358

As can be seen in figure 7-26, the l\ 14 C values for thesevarious deep water types cqrrelate with AOU and silicate remarkably well.
The lower the l\ 4 C value, the higher the AOU and silicate concentration.
This figure serves as a reminder that theconsumption of oxygen and the solution of opal act as clocks recording the relative ages of various water types.

VENTILATION OF THE DEEP ATLANTIC OCEAN
The western Atlantic Ocean receives new deep water both fromthe north and from the south, (in this regard it differs from theother deep basins which are all fed from just one source).
Aseven the mos t dense of the waters formed in the north is lessdense than the wa ter entering the western Atlantic through theVema Channel, the southern component water underrides that comingfrom the north and mixes up into it.
This mixing produces alinear 8-S array for waters deeper than about four kilometers.
These points fall very nearly along a line joining the densest
northern component water with that for the water passing throughthe Vema Channel (see figure 7-27).
The proportions of the twocontributors at any given location and depth can be obtained byscaling either one of these parameters. For example, a water witha potential temperature of 1.ooQ would be roughly a 50-50 mixtureof the two components.
Waters with potential temperatures warmer than about 2°C
yield 8-S trends with a different slope. As shown in figure 7-27,
the break between these mixing regions is qui te sharp.
We willrefer to the slope change as the Two-Degree Discontinuity (TDD).
Such a change in slope is a natural consequence of the fact that
northern component waters cover a range of temperatures and salinities.
This can be seen by returning to the one-dimensional column discussed earlier in this chapter.
Instead of considering
what happens if there is only one input at the bottom of the column, let us consider a si tuation where there are three separate
inputs.
Water with the composition of that passing through theVema Passage is injected at the bottom of the column (i.e., at 5km); water with the composition of the densest northern component
water is injected at a depth of 4 kilometers and water with thecomposition of that in the deep Labrador Sea is injected at a
depth of 3 kilometers.
The first of these waters produces an upwelling rate of x m/yr, the second produces an upwelling rate ofan addi tional lOx m/yr, and the third produces an upwelling rateof an addttional lax m/yr.
Thus, the upwelling rate beneath 4
kilometers will be x m/yr, between 4 and 3 kilometers llx m/yr andabove 3 kilometers 2lx m/yr.
In the absence of any vertical mixing the potential temperature of the water would be -0.15°C below
4 kilometers, 2. aloe between 4 and 3 kilometers depth and 2. 7 roc
above 3 kilometers depth.
However, if we introduce a moderateamount of mixing such that K/W (i. e. z*) is in all depth ranges
about 300 meters, then the temperature profile will be as shown infigure 7-2e. This will also produce a 8-S diagram similar to thatseen for the deep waters in the deep western Atlantic.
This l-D example helps explain the origin of the TDD.
If wethink of "up" as perpendicular to the isopycnal horizons, the
r

359

o

o

o~

o~

-u

U

-160

<l -190v<l -190

-220

-220

'¢

- 2 50 .......50......._""-_L-~_-.1._.....L.._ ....
80 110 140 170 200 230 260

A au (jLmollkg)

AA~CPBW
~NIDW

.~W
NPDW

- 2 50~-="~~--'-_..I-_"--.s.J~-L--Ja
60 90 I20 , 50 I80 2 10

H4 Si04 (fLmol /kg)

Figure 7-26. Plots of ~14C versus AOU (left hand panel) and versus
dissolved silicate content (right hand panel) for various deep water types.
NCW is the new water entering the western basin of thenorthern A~lantic. TDDW is water at a depth of about 4 kilometersin the western basin of the central Atlantic.
NEBBW and SEBBW
are, respectively, bottom waters from the northern and southerneastern Atlantic.
AABW is the bottom water of Antarctic origin
which penetrates into the western basin of the southern Atlantic.
CPBW is the circumpolar deep water which passes from the Pacificsector of the Antarctic to the Atlantic sector of the Antarctic
through the Drake Passage.
NIDW is deep water from the northern
Indian Ocean.
NPBW is bottom water from the North Pacific.
NPDW
is water from about 2000 meters depth in the North Pacific.
The
slope of the ~14C-AOU trend corresponds to an O 2 consumption rateof 0 .12 ~m/kg per year and that of the ~ 14C_H4SiO 4 trend to an
opal solution rate of 0.10 ~m/kg per year.

360

3.---.....~-.......- -......--..,....--,-----,8--..,,.......
U
2.....
W
0:::::> 2

~

0::

W

a..
~

w

I-

--'

WESTERNNORTH
ATLANTIC

W

12°N SloW

<t
IZ

.....
o
a. o

e VEMA CHANNEL WATER
34.7

34.8

34.9

35.0

SAL I NIT Y (% 0 )Figure 1-27. Potential temperature versus salinity diagram fordeep waters at a station in the western North Atlantic 12° northof the equator.
The prominent break in slope is designated as the
Two-Degree Dis continui ty (423).
Also shown are the compos i tiansof wa ter entering the Atlantic through the Vema Channel and ofnorthern component waters.

361

example becomes more like the actual situation in the deep Atlantic.
The
break in 8-S slope corresponds to the isopycnal ofentry of the most dense northern component water.
However, thecomposi tion of the water on this horizon does not correspond tothat of the densest northern component water.
Rather, it has acomposition matching that of the combined flux of Vema Passage
Water up through this horizon and northern component water into
the sys tern (i. e., as shown in figure 7-28 the composi tion lies
between that of the two endmembers).
From the ~14C distribution in the waters beneath the TDD, we
can get an idea as to how rapidly the deep western basin of theAtlantic is ventilated.
Were this ventilation extremely rapid,
then a plot of 14C. versus salinity should show all measurements
falling on a straight line joining the composition of the endmembers (i.e., S
34.667 0 /00 and ~14Cl = -160%0 for Vema
Channel Water and S = 34.932 0 /00 and ~ 4C = -67 0 /00 for thedensest northern component water). Actually, as seen in the upper
panel of figure 7-29, the measured values all fall below thisline, demonstrating that radiocarbon had been lost by radioactive
decay during resid nce in the mixing zone beneath the TDD.
The
devia tion of the ~1 4C from the mixing line averages abou t 20 permil.
The time required for this loss via radiodecay is about 1~0
years.
However, as we are dealing with a dynamic system thisC
deficiency is best considered within the format of the processes
ventilating this zone.
While we do not understand enough about these processes tomake a realistic 3-D model for the deep western Atlantic, we canuse the 1-D model shown infigure 7-28 to give us some idea ofthe rates involved.
Shown in the right panel of figure 7-29 arethe ~14C versus salinity trends expected for various values of the
upwelling rate produced by the input of Vema Passage Water.
As
can be seen the bes t fi t is achieved wi th a veloci ty of abou t
4m/yr.
However this fit can be obtained only if the ratio of theflux of the densest northern component water to Vema Channel Water
is reduced from the value of 10 used in the figure 7-28 example toabout 4.
The problem with this change is that it necessitates atemperature of 2.5°C and a salinity of 34.945 0 /00 for thedensest northern component water suggesting that the density ofthis water is even lower than we estimated from the properties of
the water entering the northern part of the western basin.
An upwelling rate of 4m/yr for the Vema Channel Water thenyields an upwelling rate of 16 m/yr for the most dense northerncomponent water.
As the area of western basin of the Atlantic at
a depth of 4 kilometers is about 3 x 10 13 square meters the fl~xof most dense northern component water comes out to be 4.8 x 10 4
m3/yr or 15 Sverdrups.
The composition of this water is given intable 7-5.
To this must be added the production of lower density
northern component water.
We estimate that this contribution
would raise the flux from 15 to about 20 Sverdrups.
The eastern Atlantic is fed by water originating in thewestern equatorial Atlantic. Water with a composition corresponding to the upper part of the sub-TDD mixing zone in the westernbasin spills through the Romanche Passage (located just south ofthe equator) and through the Vema Passage (located 100 north of

362

J ~

J
"

E

j

C ,~

~

.-::r:

E

10 UNITS
8= 3.55°C
S =34.950 %0

34.7 34.8 34.9 35Dr - - SALINITY (%0)
I
I

Q.

3

::r:

.....
a..
w

I

w
0

,1\

4

K
W=O.Okm.1\

B ,~____1____IOUNITS
8 =2.25°C
S =34.932 % 0

1 UNITlZ::z=~~....;..A~.-I-8 =- 0.15 °C~.L...._.L-_..L-S = 34.666%0

~

I

0

I
I

4

.-....._---a._---, 50I
23
POTENTIAL TEMPERATURE (OC)

Figure 7-28. One - dimensional model designed to demonstrate theposs ible. origin of the Two - Degree Dis continui ty.
Waters ofdiffering composition are injected into a 5 kilometer high column
a t three different levels.
This water rises and spills over the
top of the column. The 6-S diagram produced in this way resembles
that found throughout the deep Atlantic.
The shape of thetemperature profile in the column depends on the ratio K/W.
Notethat the composition at the discontinuity in e-s slope does notcorrespond to that of input water B!
Rather it lies along theline joining the compositions of water types A and B.

363

- 50 .....--r----,r---.,--~--I"'-.,

New

TOO

t

-75

-75

o

~

-100

•

...--

WESTERN
ATLANTICBENEATHJ -100o..........

uV<l -125

THE TDD

u

~

<1
-150

-I 7 5 ~---':::::'::::""'---I~-..£---"'-~~--

34.7

34.8

34.8

SAL' N , T Y (% 0)

SA L IN' TY (%0)

Figure 7-29. In the left hand panel is shown a plot of 14 C/ C ratio(expressed as Jj.14 C) against salinity for waters below the TDD in
the western Atlantic.
The compositions of the water entering thesouthern end of this basin through the Vema Channel and of themost dense northern component water are shown by the solid
circles.
Were there no radiodecay, the Jj.14 C values for the
mixtures o~ these two water types found below the TDD should fallalong the solid line.
The deviation of the Jj.14 C values below thisline is thus the result of radiodecay.
The radiocarbon measurements were made by Ostlund and Stuiver (21,22,23).
The right hand panel shows model calculations for the simple
1-D situation shown in figure 7-28.
The individual curves are fora range of upwelling rates for AABW (gi ven in meters/year).
As
can be seen a value of 4 m/yr gives the best fit.
In these calculations the ratio of K/W is set at 300 meters; the ratio of theflux of lower NADW to that of AABW is set at 4.
Hence, for thebest fit case the upwelling rate of lower NADW (i.e., the mostdense northern component water) is 16 meters/year.

364

the equator) and floods the abyssal regions of the eastern basin.

VENTILATION OF THE DEEP PAC~IC AND INDIAN OCEANS
The abyssal regions of the Indian and Pacific Oceans are eachfed by just one major source.
This source water is in each case a
mixture of northern and southern component water.
The western
Indian Ocean is fed by the spillover of circumpolar water throughthe Atlantis Passage (32°S, 50 0 E).
The North Pacific is fed bywater from the circumpolar region passing through the Samoan Passage (12°S, 179°W).
Once in these basins the bottom waters show
gradations in composition away from the point of entry (both along
the bottom and in the vertical) as the result of cross isopycnal
mixing with less dense water.
The origin of this dilution water
is not easily pinned down.
In part, it is water that has mixed
downward from intermediate depths; in part, it is water from thecircumpolar region which has been transported into the basin along
isopycnal surfaces.
The characteristics of the water in question
are summarized in table 7-5.
The reason we must be concerned with this diluting water isthat it carries different b 14C values from those of the bottom
wa ter spilling in through the passages.
Hence, as in the Atlantic, changes in the !J.14C wi thin the basin are in part the resul tof radiodecay and in part the result of mixing.
The first step in understanding this dilution is to determine
to what extent the composi tions of these deep waters can be accounted for by mixing of northern component and southern component
water alone.
We have seen that the waters of the sub-TDD mixing
zone in the western basin of the Atlantic can be generated bymixing these two components.
Is this also the case for the otherdeep waters of the world oceans? Potential temperature is a questionable tracer in this regard.
The reasons are that northerncomponent waters show a range in temperature which is significant
compared to the temperature difference between the two components
and that geothermal heating may significantly modify deep sea
temperatures on the long time scales for the ventilation of the
deep Pacific and Indian Oceans. By contrast the ranges for Sand
for P0 4 ° for northern and southern component waters are considerably smaller than the difference between mean northern and southern component water.
Thus, it is of interest to explore their use
in this regard.
Northe.rn component water has a P0 4 ° value averaging about
0.7~/kg.
Southern component water has a P04° value of about 1.3
llm/kg.
Those for the deep and bottom waters in the ocean fallwithin this range (see table 7-5). Northern component water has asalinity close to 34.93.
Southern component water has a salinityof 34.63.
Except for the 14C minimum water in the deep North
Pacific the deep and bottom waters of the ocean have salinities
falling within this range (see table 7-5).
Let us examine the distributions of these properties in moredetail.
Sections of PO 4° throughout the deep sea are shown infigure 7-30.
To the extent that deep waters are mixtures only ofnorthern and southern component waters and to the extent that the
AOD correction coefficient of 135 is valid, the contours of P04 °365

Table 7-5. Composition of the bottom water supplying the deep eastern Atlantic, thewestern Indian and the equatorial and North Pacific oceans.
Also given are thecompositions of the bott?m waters within the interior of these basins and of the deep
waters with the lowest
4C/C ratio within these basins.
Finally, the composition ofwater which has completed the loop around the Antarctic continent and is about to reenter
the Atlantic sector via the Drake Passage (between the tip of Sou th America from thePalmer Peninsula of· Antarctica) is· gi ven.
The hydrographic and chemical measurements
were made as part of the GEOSECS program (406,424,425,426).
The radiocarbon data wereobtained in the laboratories of Minze Stuiver and Gate Ostlund (21,22,23).

Water
Type

e

°c

S

04

0/ 00

0/ 00

P0 4 °
llm/kgNCW%

ECO 2 Obs.lJ. 14 c lJ.lJ.14 C
llm/kg
0/00
0/ 00

*

lJ.lJ.14 C
0/ 00

Eastern Atlantic Ocean

CJl.)

0'>

Input
No. Bottom
No. 14 C/ C min
So. Bottom
So. 14 C/ C min

1.69
1.94
2.20
1.91
2.10

34.85834.88634.90934.87734 .. 889

45.96
45.94
45.91
45.93
45.91

0.90
0.86
0.86
0.85
0.82

66
73
73
75
80

2207
2206
2209
2218
2216

0")

-113
-115
-119
-118
-123

--

-17
-25
-29
-29
-36

--

-17
-45
-71

-2
-6
-5
-10

Western Indian Ocean
Input
Bottom (7 0 S)
14C/ C min (6 ° N)

0.52
0.84
2.02

34.70734.71634.765

46.05
46.00
45.83

1.22
1.17
1.09

13
21
35

2269
2286
2334

-160
-171
-194

-11
-34

Equatorial and North Pacific Ocean
Input
0.64 34.706
Bottom (31°N)
0.97 34.694
14 C/ C min (32°N) 1.62 '34.632

46.03
45.96
45.80

1.19
1.13
1.00

18
28
--

2270
2293
2370

-162
-205
-245

2264
2264

-163
-162

--43
-83

-24
-75

--

Antarctic
Drake Pas. Bot.
Drake Pas. Deep

0.19
1.09

34.68834.723

46.09
45.96

*Difference from input water.
**Difference from endmember mixture.

1.18
1.08

20
37

--

--

-27
-25

**

Figure 7-30. Sections of the initial phosphate concentration andof salinity in deep waters of the western Atlantic Ocean (upper
panels), in deep waters in the western Indian Ocean (middle
panels), and in deep wa te rs in the wes tern Pacific Ocean (lowe r
panels) .
The measurements on which this diagram is based weremade as part of the GEOSECB program (424,425,426).
Also shown by
the dotted lines is the 04 = 45.2%0 isopycnal horizon.

367

in these sections represent the fraction northern component water(0.8 ~/kg would correspond to 83% northern component, 0.9 to 66%,
1.0 to 50%, 1.1 to 34%, and 1.2 to 17%).
The pattern in the Atlantic (at least below 2.5 km) is consistent wi th what we havealready learned.
The bulk of the deep water in the Indian Ocean
and in the equatorial and south Pacific lies in the range 1.1 to
1.2 l.1m/kg (hence, appears to be 25 ± 10% northern component).
In
the North Pacific the PO 4 ° values show a decline reaching below
0.9 in the upper deep wa ters.
This decline reflects ei ther the
input (by vertical mixing) of intermediate depth waters formed in
the northern Pacific or a deviation in the C/P ratio of theorganic material oxidized in the deep Pacific from the Redfieldvalue of 135.
C~rresponding sections for salinity are also shown in figure7-30. The pattern is similar to that seen for P0 4 °. The salinity
grades along the path of the deep current from the high values
characterizing northern component water to values closer to thatfor southern component water.
The ambient salinity for deep andbottom waters of the Indian and Pacific Ocean again corresponds toa mixture containing 25 ± 10% northern component water.
Deepwater in the north Pacific is again anomalous, its salinity isvery low.
This water has a salini ty corresponding to that ofwaters rich in southern component water and a P0 4 ° corresponding
to that of waters rich in northern component.
It appears that
this water receives a cont ribution from the downward mixing ofintermediate water.
Another way to look at this si tua tion is to plot salini ty
against P0 4 ° (see figure 7-31).
Most deep and bottom waters fallclose to the trend expected if they were produced by mixing northern and southern component water.
Deep waters from the Pacific do
not.
They have P0 4 o-S coordinates falling well below that
producted by endmember mixing.
As shown in figure 7-31 plots of potential temperature versus
PO 4 ° for the same wa ter types fall wi th one exception (i. e., Indian Ocean 14 C/ C. minimum water) within the composition triangle
for mixtures of northern and southern component water.
Suprisingly the si tuation for 8 and P0 4 ° is more concordant than thatfor Sand P0 4 ° (see figure 7-31).
A possible reason for this isthat the 8-P04° coordinate for the third contributor to North
Pacific deep water lies along the line joining the coordinates ofnorthern and southern component water.
This just demonstrates how
difficult this ~roblem is.
If we are to sort it out we need highly accurate c 1(or cD) data for these waters.
Like salinity
these isotopic compositions are truly conservative.
Unfortunately, samples collected for this purpose as part of the GEOSECS
program were never analyzed.

°

THE GRAND CYCLE OF RADIOCARBON IN THE DEEP OCEAN

The mean Do 14 C value for wa ters at depths grea ter than 1. 5
kilometers is about -1750/~o.
The mean EC02 concentrati0{7 f~rthese waters is 2.3 moles/m.
Their volume is about ~ x 10m.
Hence, the deep sea contains about 0.825 x 1.20 x 10- 2 x 2.3 x 8x 10 17 or 1.8 x 10 6 moles of radiocarbon.
One part in 8200 or

368

0

~
0

-u 2

34.9

.......

0.......

>~- 34.8
z

-...J
~

U)

Q..•

~

w
~

•

34.7

.-:

ATLANTIC~OA INDIAN
8 PACIFICo ANTARCTIC
0.9

0.7

1.1

-I

1.3

0.7

P04° (p.mol/kg)

3- 34.9
->-

-u 2

0

~
0

0
.......
•
0:E..... 34.8

-

Z

-...J

«
(J) 34.7w~4CIC
MINIMUM

•

~

~

0

a.

WATER

0.7

14 C/ C

~

0.9

r. I

1.3

P04° (p.mol/kg)°-I

MINIMUM

WATER

0.7

0.9
I. I
P04 (,arnol/kg)

1.3

Figure 7-31. Plots of salinity and potential temperature against
initial phosphate concentration for the major bottom waters (upper
panel) and for the maj or 14 C/ C minimum
waters (lower panel) in
the world ocean.
The compositions of the endmember waters arealso shown.

369

about 222 moles of these 14C atoms undergo decay each year.
The
14
11
C value for the two major sources of ventilation of the deepsea are -67 0 /00 (for northern component water) and -154 0 /00(for southern component water).
The entry of this water and therain of carbon-bearing particulates must balance the decay ofradiocarbon occurring in the deep sea.
Northern component wa terenters the deep sea at the rate of about 6 x 10 14 m 3/ yr (i.e.,
about 20 Sverdrups).
It carries about 2.1 moles/m 3 of Ico 2; its
14C to C excess over ambient deep water is (0.933-0.825) 1.2 x
10- 12 or 1.3 x 10- 13 atoms 14 C per atom C.
Thus, northern comronent water carries with it about 1.3 x 10- 13 x 2.1 x 6 x 10 1 or
164 moles of 14 C . Particulate debris rains to the deep sea carrying carbon wi th a 11 14 C value of about -70 0 /00.
The rate of fallof such debris is about 0.5 moles/m 2 yr.
The area of the ocean at
the 1500 meter contour is about 3 x 10 14 m2 •Thus about 1.5 x
10 14 moles of carbon reach the deep sea each year in part icula teform.
Almost all of this carbon is returned to solution.
The
excess 14C!~ ratio for this carbon is (0.930-0.825) 1.2 x 10- 12 or
1.3 x 10- 1 .
Thus, raining particles carry about 19 moles of
14C•
Combined wi th the 164 moles supplied by northern component
water we have now explained the resupply of 183 moles out of the
222 lost to radioactive decay.
The remaining 39 moles of radiocarbon presumably accompany
new southern component water.
As there is no well-confined miXing
chamber into which this water flows\ an estimate for the flux ofthis water cannot be made from the 1 C distribution, as it was forthe northern component water.
One approach is to consider the
relati ve proportions of northern and southern component water inthe deep sea. Based on P0 4° data, it appears to be about 0.25 New
and 0.75 SCW.
If the mean residence time of the two components
in the deep sea were similar then three times as much southerncomponent water mus t reach the deep sea each year than northerncomponent water.
However, as the SCW is the most dense water inthe sea and flows along the bottom of the ocean it is probably replaced on a longer time scale than NCW which overrides its more
dense counterpart.
We instead estimate its flux to be about thesame as that of NCW (i.e., about 20 Sverdrups).
The flux of southern component water would then be 6 x 10 14
3
m /yr. The ~C02 content of this water is about 2.2 moles/m 3. Its
excess 14C/ C ratio is (0.846-0.8251 1.2 x 10- 12 or 2.5 x 10- 14
atoms 14C per atom e.
The flux of 4e into the deep sea via thissource is hence 2.5 x 10- 14 x 2.2 x 6 x 10 14 or 33 moles/yr.
These calcu'la tions are summarized in table 7-6. Roughly 220 molesof 14 e are replaced in the deep sea each year.
These considerations sugpest that the total flux of water
into the· deep sea is 12 x 10 4 m3 /yr ~i. e., 40 Sverdrups).
The
volume of the deep sea is about 8 x 10 1 m 3 • Thus the characteristic flushing time is about 670 years.
While this result is uncertain due to an inadequate knowledge of the rate of formation ofsouthern component water, it is significantly less than the 1000
year figure given in chapter 3.
Why the difference? Part of thereason lies in the fact that we have considered only the flushingof the sea below 1500 meters. Had we included the entire deep sea
the time would have to be increased.
The other reason is that in370

Table 7-6. Summary of information needed to examine the budget of
14 C in the deep sea.
Loss
Via radioactive decay
Volume of water
Mean 2:q02
Mean !J 4C
Mean 14 C/ C
Amount of 14 C
Amount 14 C decaying indeep sea (below 1.5km)-8 x 10 17 m3
2.3 moles/m 3-175 0 /00
1.00 x 1~_12
1.8 x 10 moles
222 moles/yr

Gain
Via northern component waterFlux-6 x 1014m3/~r
2.1 moles/m-67%00.13 x 10- 12
164 moles/yr

2:~O 2!J 4 C

14 C/ C - 14 C/ Cncmean deep sea
Flux of 14 C
Via southern component water
Flux-6 X 1014m3/~r~CO 2

2.2 moles/m-154 0 /00
0 .. 025 X 10- 12
33 moles/yr!J14 C

14 C/C - 14 C/ C
mean deep sea
sc
Flux of 14 C
Via particulate rain
Carbon flux***!J14 C

14
14 /C
C/Cparticles- C mean deep sea
Flux of 14 C
Total flUX of new 14 C to deep sea

0.5 moles/m 2yr-70%00.126 x 10- 12
19 rnoles/yr
216 moles/yr*The mean preformed P0 4 content of northern component water is
0.711m/kg and that of southern component water is 1.3lJm/kg.
Ifthese two water types are fed to the deep sea in nearly the sameamount then the preformed P04 content of new deep water ave rages
1.0llffi/kg.
The P04 content of deep water averages about 2.3lilll/kg.
Thus abou t 1.3lJm/kg was received by particles.
If the C/P ratio
for these particles is 125 then about 160 llm/kg or about 7% of the
carbon was. sUPPtie~ by particles.
The ipput of carbon Ro the deepsea is 12 x 10 4m /yr x 2.15 moles C/m or 2.6 x 10 lmoles/yr.
Seven percent of this is 1.8 x 10 14 moles/yr.
The area of the sea
at 1 .. 5 km depth is about 3.2 x 10 14 m2 "Hence the flux is 0.56
moles/m 2yr.
**Because of the 20%0 depletion in 13 C relative t9. 12 C during
photosynthesis there must be a 40 % 0 depletion in
4C relativet~ 12C.
For CaC03 there is no depletion in 13 C.
Hence the mean
!J 4C value for the raining organic rna t te r should be -80 0 /00 andt~~t in the raining CaC03 -40%0.
The composite should have a!JC value close to -70 0 /00.

371

the calculation in chapter 3 we assumed that the water ventilating
the deep sea carried a A 4 C of -50 0 /00.
As we have seen the actual values are far lower (-67 % 0 for northern component waterand -1540/00 for southern component water and hence -125%0
for a 1-to-2 mixture of these waters).
These l4C-based rates are consistent with current meter results from key areas in the ocean where strong confined flows occur. Based on results from the northwestern Atlantic, Worthington
of Woods Hole Oceanographic (701) estimates a flux of about 12
Sverdrups of new North Atlantic Deep Water (i.e., of northern component wa ter) .
Nelson Hogg of Woods Hole Oceanographic placed
current meters in the Vema Passage and established a flux of 4
Sverdrups of Antarctic Bottom Water (i.e., of southern component
water) into the central western basin of the Atlantic (701).
Current meter resul ts from the Drake Passage reveal that deep
water is being recirculated through the Antarctic at a rate of 130
Sverdrups (703,704).

BIOLOGICAL SHORT-CIRCUITING
As if we do not already have enough complications in our attempt to decipher the distribution of l4 C in the deep sea there
are two more which must be considered. First, particles carrying
l4C~ to the deep sea short-circuit the mixing cycle.
The apparent
ages we get f rom 14C~ reflect a compos i te between the age of the
carbon carried by the water and the carbon carried by the particles.
Second, even prior to the era of nuclear testing the l4 C/ C·
ratio in the atmosphere varied wi th time.
Hence, the carbon in
the deep sea need not have left the surface with the A l4C values
we reconstruct for the time immediately preceding the nuclear era.
In table 7-7 we give the differences between the Al4C value
for deep waters within a given basin and the A 14 C for the bottom
wa ter feeding these basins.
We also give the l:CO 2 contents ofthese waters.
If we attribute the increase in l:C02 to the destruction of particles within the basin, and assume that theseparticles originated wi th the A l4 C for temperate and equatorialsurface waters (i. e., averaging -70%0) we can then correct forthe contribution of particles to the l4C/C ratio.
The magnitudeof these corrections can be determined through a qui te simple
calculation.
We will treat each water type shown in table 7-7 asa well-mixed reservoir fed directly by the input water. In such asituation the carbon balance is given by:
7-6
where B is the flux of particulate carbon into the reservoir, R is
the flux of input water into the reservoir, and [l:C02]r and[l:C02]B are the concentrations of dissolved inorganic carbon in
the input and in the basin water, respectively. The corresponding
material balance equation for radiocarbon is:
B 14 C/ C )p + R[l:C02]r 14 C/ C)r = (R+AV) [l: C0 2]B 14 C/ C )B
whe re14C/ C) r,

14C/ C )B ,and

14 C/ C ) pare ,

372

7-7

respectively,theTable 7-7. Comparison of ventilation times calculated disregarding the contribution ofparticulates (T) with those calculated including the contribution of particulates (T*).
Water
Type

C-A,)

[~CO 2] Input\lm/kg[~CO 2 ]Basin\lm/kg

14

1~/C

)

InRut
C/C)atm

T*Basinatm

1'J3asinyr

14 C/ C )Basin

14 C/ Cyr

No. East
Atl. Bot.

2207

2206

0.887

0.885

18

18

No. East
Atl. 14C Min.

2207

2209

0.887

0.881

56

56

So. East
Atl. Bot.

2207

2218

0.887

0.882

46

49

So. East
Atl. Bot.

2207

2216

0.887

0.877

94

95

WesternInd. Bot.

2269

2286

0.840

0.829

109

115

WesternInd. 14C Min.

2269

2334

0.840

0.806

346

371

WesternPac. Bot.

2270

2293

0.838

0.795

444

453

WesternPac. 14C Min.

2270

2370

0.838

0.755

901

944

.........

C-A,)

radiocarbon to stable carbon ratios in the input water, the basinwater, and in particulate matter, and A is the decay constant for
l4 C.
These equations can be combined to eliminate B.
Also
T*basin, the basin ventilation time, can be substituted for theratio VIR and T l4 C, the mean life of l4 C, can be substituted forthe l/A. The resulting equation is:
T*basin --7-8

l4 C/ C )B

14 C/ C)atm

This may be compared to the ventilation time, Tbasin' obtained
if the particulate contribution is neglected (i.e., if [LC02]I =[ LC0 2]B):

Tbasin=7-9
l4 C/ C )B
l4 C/ C)atm

The results of such comparison are shown in table 7-7.
While theabsolute values of the residence times are not to be taken as
highly meaningful (the configuration of the ventilation model is
not adequate), the ratio of one time to the other is meaningful.
In other words a more realistic model will give different absolutetimes but not different ratios of the times.
The correction forparticulate rain will be on the order of 10% or less. This is thesame result as was obtained for the deep ocean as a whole (19 of
the 216 moles of llf C entering the deep sea come with particulates). Thus, while a correction should be made, the error introduced by its neglect lies well wi thin other uncertainties of theabsolute ages.
TEMPORAL VARIATIONS IN RADIOCARBON PRODUCTION

The second complication stems from the fact that even before
the onset of nuclear testing the l4 C/ C ratio in the atmosphericcarbon did not remain cons tant.
This has been shown by precise
l4 C/ C measurements on tree rings.
The record reconstructed byMinze Stui ver at the Uni versi ty of Washington for the last one
thousand years is shown in figure 7-32.
The results have beencorrected for radioactive decay since growth, using the chron~~ogy
obtained by counts of the growth rings.
As can be seen the
C/Cratio in wood grown immediately before the onset of nuclear test-374

---r--.. . . --.----.0111......

-~-..._- ..

20 r - - - r - - - r - -........
MEAN
MILLENNIAL

_VA

LUj!_

ATM
-10

-20
- 30
1000

......_ ......._ - - L .__ . L _ - . l_ _I . . - _ . a _ . _ . . . I _ _...

~_

1200

1400

1600

leoo

2000

YEAR
Figure 7-32. 14 C/C ratio (expressed as f).14 C ) for atmospheric CO 2
over the last 1000 years as reconstructed from measurements madeon tree rings.
The fluctuations prior to 1850 are thought to be
the resul ts of changes in the acti vi ty of the sun.
The decline
over the last 100 years is mainly the result of the addi tion of
14 C free
CO 2 to the atmosphere through the burning of fossil
fuels.
These measurements were made by Minze Stui ver of the
University of Washington (19,26).

375

40 .....- ......- - . - -........- - r - - - - r - -........- - - .

20
0
~

0

0~

~

vU
<J

-20

ATMOSHERIC CO 2 (TREES)SURFACE OCEAN IC0 2 (CORALS)

-40
-60

-80 .....- .....- ......----.--.&..-----------......

1600 1650 1700 1750 1800 1850 1900 1950
DATE OF GROWTH (yrs)

Figure 7-33- Comparison of the 11 14 C record f?r the atmosphere(based on measurements on tree rings) and the 11 4 C record for thesurface ocean (based on measurements of coral rings) for the 1640
to 1900 AD time period.
The coral measurements were made by Ellen.Druffel in the laboratory of Hans Suess at the University ofCalifornia at San Diego (15,16).

376

ing was lower than at any time during the last millennium.
Threecauses for these variations have been identified.
First, there
has been a long-term change in the strength of the earth I s magnetic field.
Since this field acts as a partial shield against
the incoming cosmic ray protons\ variations in its strength leadto varia tions in the rate of 1 C production.
These variations
lead to changes in 14C production on the time scale of several
thousand years.
Second, there are shorter-term variations in
condi tions on the sun's surface (visible as "sun spots") whichalter the sun's magnetic field.
These changes also modulate thera te of 14 C. production in the earth's atmosphere.
They are responsible for the peaks in 14 C/ C ratio spaced at rou~hly 200-year
intervals.
Finally, the emission of C02 free of
4 C from theburning of fossil fuels has caused the post 1850 decline in 14C/ C
ratio.
Because of this, the 14 C/ C ratio in the atmosphere wasabout 22 0 /00 lower during the early 1950' s than it was on the
average during the last millennium.
These changes must certainly have been reflected in change s
in the 11 14 C value for the sources of bottom water.
However, because of the complexity of the formation process it is not easy togauge the relationship between the atmospheric and ocean chang s.
It is clear, however, that the changes in the source wa ter 1114 C
values will not be as large as the changes in atmospheric 11 14 C
values.
Direct evidence is available to demonstrate that this is
the case.
The coral record obtained by Druffel at U. C. San Diego
for the Florida Straits extends back to the year 1650.
As can beseen from figure 7-33, the magnitude of the changes in the surface
ocean 14 C/ C ratio is much less than for the atmosphere.
For
example, there is no equivalent of the strong maximum in atmospheric 14 C/ C ratio close to 1700 AD in the oceanic record.
Between 1650 and 1900 AD the 11 14 C value for water in the Caribbean
and Florida Straits remained constant to within ±5 0 /oo.
Figure7-34 shows the data on which the surface ocean curve is based.
Since the scatter is quite large, the validity of the small
wiggles seen in the mean curve is open to question.
In thisfigure the record from 1900 to 1950 is included.
Here a definite
decrease paralleling that for the atmosphere is seen.
Itsmagnitude is only about 8 per mil, which is 40 percentof thatobserved for the atmosphere over this period.
Thus, whilevariations in the production rate of 14 C must be considered inconnection with the interpretation of the 14 C distribution within
the sea, in most cases their inclusion will produce only smallchanges in the results.
ARGON-39

One of the radiotracers, 39 Ar , mentioned in chapter 3, has
yet to be discussed.
As 39 Ar is produced by cosmic ray impacts
and has a half-life of 280 years, its steady state distribution
wi thin the ocean should have much to say about the time scale ofdeep sea ventilation. Not only is its half-life more suitable forthe task than that of 14 C , but unlike 14 C the 39 Ar content of new
deep water should be at eqUilibrium with the atmosphere. Our reason for excluding this important isotope from our discussion of377

- 20 r - r - - -.......----r---...,...--....,....---..----.--..

...........
0

--40

~
0

-.",

q-U

-

<J -60- 80 ....- 1650
1700

...-.
1750

1800

......._ _- & ._ _- - 1 1 - . - - - 1
1850
1900
1950

YEAR

Figure 7-34. 14 C/ C ratios (expressed as ~14C) on growth ring dated
coral samples from the Florida Straits and from the coast of
Belize (Honduras) from 1650 to 1950. The post-1900 drop resulting
from the input of fossil fuel C02 is apparent.
The measurements
were made by Ellen Druffel in the laboratory of Hans Suess at the
University of California at San Diego (15,16).

378

deep sea ventilation is very simple:
it has yet to be measured in
the deep sea.
The reas on is that its abundance is so low, tha teven when the world's best low background counters are employed todetect its acti vi ty, about one ton of water must be processd inorder to get enough 39 Ar to measure!
Currently only the laboratory at the Uni versi ty of Bern in Swi tzerland has the capabili ty
of counting the 39 Ar obtained from even these very large water
samples.
At the time this book was being written several groups
were collecting larger water samples and stripping the gas fromthem for delivery to Dr. Loosli in Bern.
Hence, in a few years,
there should be some new things to say about the time scales ofdeep ocean ventilation.
When considered together, the dis tributions of 39 Ar and 14 C will put much stonger constraints on ventilation models than is possible with 14 C alone.
SUMMARY

The deep sea is ventilated from two small areas on the ocean
surface.
Northern component water is generated at the north endof the Atlantic, and southern component water is generated in the
Weddell Sea on the perimeter of the Antarctic continent. Althoughthese two water types have quite different temperatures, salinities and initial phosphate contents, their contributions to thewater masses in the deep sea can be estimated in only an approximate manner.
The paths by which these waters travel through the
deep sea, although known in a general way, are not understood in
detail.
Diffusive mixing in the horizontal and in the verticalgreatly blur the advective component of these motions.
The time scales for movement of these isotopes through the
deep sea are largely based on the distribution of radioactive
tracers.
Shortlived 222 Rn and 228 Ra offer information regarding
vertical and horizontal diffusive mixing within a single deep
basin.
14e provides the cu rrent basis for es tima ting the timescale for the ventilation of individual basins.
However, due tothe small extent of atmospheric equilibration for the carbon in
the waters sinking from the surface of the Weddell Sea, the radiocarbon method does not give us a complete answer. The hope of the
future is that 39 Ar data will allow this gap to be closed.
PROBLEMS

7-1 The following results were obtained formeasurements in near-bottom waters:aprofileDistance above
Bottom, meters

Radon Activity,
dpm/lOO liters

200
100

20.3

50
40

20.7
21.2
24.1
36.3

19.8

30
20
10

84.0

379

of

radon

What concentration of radon is produced by the 226 Ra dissolved in the water column? What is the apparent coefficientof vertical eddy diffusivity as determined from the distribution of excess (sediment-derived) radon?
7-2

If the coefficient for vertical (i.e., diapycnal) mixing
averages 1 cm 2/sec in the deep sea, then what must the coefficient for horizontal (i. e., isopycn~l) mixing be in order
to explain the distribution of excess He in the deep Pacific
Ocean?

7-3

How great must the coefficient of horizontal mixing b:f2S inorder to be able to explain the near-uniformi ty ofRaalong isopycnals in the deep Atlantic?

7-4

What would the ratio of the 14 C/ C in planktonic forams tothat in benthic forams be in material formed recently in the
North Pacific Ocean?
Assuming that the rate of deep sea
ventilation has remained constant with time what would be theratio for material formed during the peak of Holocene time(i.e., about 6000 years ago) and for material formed during
the peak of glacial time (i. e., about 18,000 years ago)?
Were this method adopted to determine the ventilation rate of
the ocean during glacial time, what size samples would beneeded if an accuracy in the ventilation time of 200 years
were desired?
The conventional 14 C method involves themeasurement of the. beta particles emitted when 14 C atoms undergo radiodecay.
Assume that these betas could be detectedwith 100% efficiency and that there was no background for thecounter.
Counts of one month would be made on each sample.
A new method under development at the time this book was being wri tten invol ved measuring indi vidual 14 C atoms in ahigh-energy mass spectrometer.
Assuming a detection efficiency of one a tom in 10 3 in the sample used and again zerobackground, what sample size would be needed? Benthic forams
weigh about 3xl0- 5 g each.
The coarse fraction (>200 microns) in a deep sea core cons is ts almos t entirely of foramshells.
Typically, about 3% are benthic (i.e., bottomdwelling) and the rest are planktonic (i.e., surface-waterdwelling) .
How many forams would have to be viewed to findthe required weight of benthics?
How many benthics would
have to be picked?
If a person could find and isolate one
benthic foram each minute, how many man hours would be
required to prepare a sample?

7-5

Assume that deep waters form with an equilibrium amount of
atmospheric gas.
If so, what would you predict the ratios ofthe specific activity of 39 Ar in argon extracted from thefollowing key water types to that in atmospheric argon to be?
Two-Degree Discontinuity Water
Southeastern Atlantic l~C Minimum Water
North Pacific Bottom Water
North Pacific 14 C Minimum Water
Drake Passage Water

380

What are the major ~roblems connected with making such estimates?
Would the9Ar/ Ar ratio in new southern component
water as defined here be equal to that for the atmosphere?
Why?
7-6

Were communication between the deep sea and the thermocline(plus surface ocean and atmosphere) suddenly cut, at what
rate (% per hundred years), would the 14 C/ C ratio in mean
deep sea carbon and in atmosphere + surface ocean + thermocline carbon change?
Assume that ninety percent of the
carbon lies in the deep sea.
Might climate-induced changes
in the rate of deep sea ventilation then provide an alternate
explanation for the variations in atmospheric 14 C/ C ratio
reconstructed from tree ring measurements?

SUPERPROBLEM#7

A geochemist named Gregg Hormone becomes inr~rested in therelationship between the distributions of 14 C ,Ar and excess
mantle derived 3He in the ocean. He extends the computer aquariumof Stover and Eastland to include these two noble gases.
Concerned about the role of the Antarctic Ocean, he decides to modify
the number of boxes and their configuration.
The box configuration and water fluxes he adopts are shown in figure 7-35.
Hormone decides, for simplicity, to neglect the biologicalcycles in the sea and assigns all the reservoirs a EC02 concentration of 2.2x10- 3 moles/liter.
He generates 14 C and 39 Ar atoms
in the atmosphere and adds 3He to the Indopac reservoir (but not
to Atlantis).
Gas exchange between the atmosphere and ocean isassumed to be limited by a stagnant film 50 microns in thickness.
The areas of the interfaces are gi ven along wi th the volumes of
the reser~oirs in figure 7-35.
The area of the Antarctic outcrop
is kept small in order to approximate the impedance of the thinlayer of low salini ty water which covers much of the Antarctic
region.
Hormone then turns on the computer and allows the distributions of 14 C , 39 Ar and 3He to come to steady state.
What does hefind for the 14 C/ C ra tio in each of the fi ve reservoirs comparedto that in the atmosphere?
Compare the 14 C{C ratios with thoseactually observed. What is the ratio of the He excess in each of
the reservoirs relative to that in the Indopac reservoir? The observed 3He , excess for the Antarctic reservoir is about 0.4 that in
the deep Pacific, and the 3He excess in the Atlantic is about 0.1
that in Indopac. How do these compare with the computer model results?
What does Hormone find for the 39 Ar/ Ar ratio in each of
the reservoirs compared to that in the atmosRhere?
How does theapparent radiocarbon age computed from the 39 Ar/ Ar depletion in
Indopac water (relative to the atmospheric argon) compare With theapparent radiocarbon age compu ted from the ra tio of the 1 C/C depletion in Indopac water compared to warm surface water? What implications does this have to the use of 3He to determine the rateof hydrothermal circulation in the ridge crests (see chapter 6)?

381

-+-

J

-

112

.. I

"'

24

16 1...
1-

~

....

~ 8
.,

I NDOPACANT.

ATLANTIS

r

J~

8

20

....

1'."'"
I

4

---

32

-

VATLANTIS= 2.7 x 10 17 m 3

VATL. SURF.

= 0.3 x 10 '7 m3

V,NDOPAC= 8.1 x IO'7 m 3

VrNPAC. SURF. = 0.9 x 10'7m 3'7 3
VANTARCTIC = 1.5 x 10 m
A ATL. SURF.

= 0.9 x 10'4 m 2

A I NPAC. SURF.

= 2.7 x 10'4 m2

AANTARCTIC = O.02x IO'4 m 2

Fi.gure 7-35. C.onfiguration, fluxes, volumes
for the reservoirs in the superproblem.
Sverdrups (i.e., 10 6 meter 3 /second).

and interface areas
The fluxes are in

LAST MINUTE ADDITION (SEPTEMBER 1982)
Loosli of Bern, Switzerland, has conducted several 39 Ar measurements on argon extracted from North Atlantic Deep Water by
Smethie of Lamont-Doherty on the TTO expedition and by Roether ofHeidelberg, West Germany on the METEOR expedition (both in 1981).
These as yet unpublished results show a wide range of 39 Ar values,
which are consistent with the accepted pattern of ventilation andthe rough time scale provided by radiocarbon.
Thus not only has
Loosli demonstrated his ability to carry out reliably this very
difficult measurement of the radioactivity of argon extracted fromsea water, but also has proven its power as a sensitive time
tracer for deep sea ventilation.

382

Ckapte~ i
THE ANTHROPOGENIC INVASION
THE MOVEMENT OF WATER THROUGH THE OCEANIC THERMOCLINE

INTRODUCTION
The activities of man result in the release of a host of substances to the environment.
As might be expected, much of thismaterial eventually reaches the sea.
For our purposes we will
divide these substances into two ma.jor categories.
In the firstcategory are those whose bulk is sufficient to impact the sea's
chemistry.
The lead released during the combustion of gasoline,
the CD2 generated by the combustion of fossil fuels, the oxides ofnitrogen produced when nitrogen combines with oxygen in automobile
engines, the phosphate mined for use in fertilizer and detergentsand the S02 created during the combustion of coal and oil fall into this group.
In the second category are those trace substances
which can serve as tracers of oceanic processes.
The 1 4 C , 3 H and
90 Sr produced during testing of nuclear devices, the freons lostfrom refrigerator coils and spewed forth from spray cans, and the
85 Kr produced in nuclear reactors fall in this group.
In addition
there are, of course, substances which have not yet been implicated with regard to possible chemical impacts and for which
tracer applications have yet to be found.
In this chapter we focus our attention on the anthropogenic
tracers.
They constitute the most powerful tool we have for the
study of the ventilation of the main oceanic thermocline.
It is
as if man were conducting an elaborate dye experiment to see how
substances added to the ocean surface are carried to its interior.
Through a knowledge of the patterns and time histories of the
input of these tracers to the sea surface and a knowledge of theevolution of their distributions within the sea, much insight canbe obtain~d with regard to the patterns and rates of circulationof waters in the ocean's upper kilometer.
INPUT FUNCTIONS

Although added to the environment over roughly the same timeinterval, the information obtained from individual anthropogenic
tracers is not totally overlapping.
Each tracer has its own special story to tell.
To understand why this is the case we must
look into the differences in the inputs of these tracers to thesea.
One characteristic of interest in this regard is the equilibrium partition coefficient of the tracer between air and sea

383

(i. e., the ra tio of the amount of the tracer in the sea to theamount in the air after steady state partitioning has beenachieved).
We define the partition coefficient as the ratio of
the amount of tracer in the sea to that in the air.
For tritium
and radiostrontium the partition coefficients are very large.
Virtually all the tracer will be transferred from air to sea. For
radiokrypton and the freons the distribution coefficients arequi te small.
Over 90% of these gases will remain in the atmosphere.
Radiocarbon falls in between these extremes.
At steadystate about 1.5% of the 14 C will reside in the atmosphere.
The
differences among these three groups of tracers becomes even more
pronounced when their distribution in the year 1973 (the mid-pointof the GEOSECB surveys) is considered.
At that time the ratio ofanthropogenic 3H in the ocean to that in the atmosphere was about
1000, for anthropogenic 14C it was about 1, and for anthropogenic
85 Kr it was about 0.01.
The fact that the wind-mixed upper layerof the sea can contain all the 90 Sr and 3 H reaching it from theair leads to a very rapid transfer of these isotopes from thetroposphere to the sea.
These differences among the tracers affect theirdistributions in the sea in two major ways.
First, they produce differences in the geographic pattern of entry of the tracers into thesea.
Due to their very short time of residence in the atmosphere,
3H and 90 Sr reach the sea surface in the same latitude belt as
they were added to the troposphere.
As far more nuclear testing
was carried out in the northern than in the southern hemisphere,
most of the tritium and radiostrontium reside in the oceans of thenorthern hemisphere.
By contrast, 14 C, 85 Kr and the freons had
spread throughout the atmos~here before entry into the sea.
Hence, the inputs of 14 C and 5 Kr to the sea are nearly symmetrical about the equator.
A second way in which the contrast in partition coefficient
imprints itself on the distributions of these substances in thesea is through the differences in the time histories of their inputs to the sea.
The 3 H and 90 Sr atoms generated during nuclear
tests reached the sea surface wi thin a few weeks after their injection to the troposphere.
The debris from the large tests(which gene ra ted mos t of the radioacti vi ty) was carried throughthe troposphere into the stra tosphere.
Each spring stra tospheric
air is injected into the troposphere at lati tudes of about 45°.
These injections brought down the debris from the stratosphere to
the tropos.phere in a pe riod of a few years.
The tes ts of verylarge nuclear devices were conducted by the USA, UK and USSR justprior to the signing of the test ban treaty in 1962. Most of the"fallout" from these tests reached the troposphere during 1962,
1963 and 1964.
Thus at the time of the GEOSECS surveys (19721977), these isotopes had been in the ocean ror about ten years.
By contrast, at the time of these surveys, the 14C produced during
these same tests was still entering the sea in sizable amounts.
As we shall see, this difference in input histories has particularimportance to our understanding of processes taking place in thethermocline of the equatorial ocean.
As mentioned above tritium is delivered to the sea surface in
two ways: THO molecules are incorporated in rain drops which fall384

onto the sea and they impact the sea surface as water vapor.
The
contribution of rain has been assessed through measurements of therate of rainfall and of the tritium content of rainwater duringthe "fallout" years. The time dependence of the tritium concentration in rain at a single station is shown in figure 8-1 and thela ti tude dependence of the t ri tium content of rain is shown in~igure 8-2.
This entry mechanism explains only about 20% of the
H present in the ocean.
The remainder must have entered throughthe impact of water molecules onto the sea surface (a process far
less easily quantified than delivery by rainfall). The geographic
and time patterns of tri tiuro input to the sea via vapor impact
must be reconstructed in a less direct way.
The T/H ratio inwater vapor is assumed to be proportional to that in rain.
As
rains (but not vapors) from many places on the earth's surface
were collected and analyzed for tritium during the critical years,
this assumption prOVides an index of the delivery rate of tritium
to the sea surface via vapor impact at any given place and time.
The absolute delivery rate is then obtained by forcing the integrated ocean-wide vapor impact to match the amount of bomb-produced tritium found within the sea minus that accounted for by
rainfall. *The resul ts of the 3 H input reconstruction carriedout by Roether and Weiss are summarized in table 8-1.
Two additional points must be made here with regard to tritium.
First, the contribution of tritium produced by cosmic ray
interactions is so small compared to that of tritium produced by
nuclear testing that all the tritium measured during the GEOSECS
surveys can be treated as anthropogenic.
Second, the fact that
tritium is radioactive and disappears with half-life of 12.2 years
is not important to its use as an anthropogenic tracer.
Its utility would be exactly the same were it stable.
In fact, in thediscussions which follow we will often normalize the tritium concentration for seawater samples collected in different years to
the amount -of tri tium they would have contained hadradiodecay
not occurred.
What we have just said for tritium applies equally to 90 Sr .
There is one difference:that 9 °Sr' s reaching the continents inrain or by aerosol impact was absorbed very efficiently by soils.
This has two implications.
First of all, this 90 Sr does not movefrom continent to sea as does 3 H.
Secondly, soils from flat
grass-cover~~ land
provide excellent integrated records of She
amount ofSr fallou t.
Thus the la ti tude-dependence of 9 Sr
fallout can be reconstructed directly from soil data.
The recordof 90 Sr delI very as function of time and of lati tude are also
shown in figures 8-1 and 8-2.
As we see from these diagrams, thefallout patterns of tritium and 90 Sr are remarkably similar - sosimilar, in fact, that one might conclude that the information
90 Sr yields about the ventilation of the sea is no different from
*Actually, about 40% of the tritium added to atmosphere initiallyreached the continents rather than the ocean. Most of the tritium
which reached the continents was subsequently transferred to thesea by either river runoff or by evaporation followed by molecularimpact or rainfall on the ocean. This complicates the reconstruction of the input pattern and input history of tritium to the sea.

385

Figure 8-1. Plot of the mean 3 H content of rain at Valencia, Ire-

land from 1952 to 1974 (solid circles).
Also given is the total
annual northern hemisphere 90 Sr deposition (open circles).
Fromthis comparison it is clear that the time history of the input ofthese two isotopes is quite similar. The differences are related
to the ratio of escaping neutrons (producing 3H) to uranium fissions (producing 90 Sr ) for the various bombs tested. This diagram
was pUblished by Dreisigacker and Roether (61).
12
30

10

-

20 .-

8

U

:3

+-

It)

(\J

o

-

-

15 0

6

4

o
I

2

,0--0- -0- -0

"

,9?,
, ,,

",
\

,\

I

5

I
I

I

YEAR
Table 8-1. Summary of the amounts of tritium added to the major
oceans as of Jan. 1, 1973.
The contribution of river inflow isseparated from that of precipitation and vapor exchange.
The
units are 10 6 curies. This compilation was done by Weiss andRoether of Heidelberg, W. Germany (65).
Northern Hemisphere

Southern Hemisphere

Rain +
Rivers Total
Yap. exch.

Rain +
Rivers Total
Yap. exch.

N

+ S

Total

Atl.

520

175

695

105

10

115

810

Ind.

50

40

90

160

5

165

255

Pac.

650

60

710

230

10

240

950

A+I+P

1220

275

1495

495

25

520

2015

386

35

-E

N

30

~

........

N
25 E

u
rt1

~
en

10

Z
0

E
0
.-

20

-r-

0

~

-

15 0

U')

0

l0-

a..

en

w

10 0

en

0
l0-

U')

0

/0

5

en
0

30

60
40

6
35

5
E

~
en
25 E
0
..20 0

.- 4
~

....,
0

-:r:

-

N

30 E

en

3

0

/5

rt)

~

:r:

2

10 ....,ATLANTIC .....
INDIANPACIFIC

0

LATITUDE

30

60

5
0

N

Figure 8-2. The upper panel shows the amount of 90 Sr accumulatedon the earth's surface up to the year 1968 as a function oflati tude.
These estimates are based on the analyses of soils.
The lower panel shows the latitude dependence of the total amountof 3H reaching a unit area of ocean surface for the three major
oceans as of 1973 as calculated by Weiss and Roether (65).

387that yielded by 3 H.
While this is true we shall soon see, however,that there is a good reason to consider both theseisotopes.
The radiocarbon produced during nuclear testing enters theocean by molecular diffusion through the stagnant boundary layer.
We get its input to the sea at any given place through a knowledgeof the CO 2 exchange ra te between ai r and sea and of the temporal
records of the 14 C/ C ratio in atmospheric C02 and in surf~ce ocean
EC02"
The latter is needed because some of the bomb
4 C atoms
which enter the sea come back out again.
Records for three major
lati tude bands are shown in figure 8-3.
The atmospheric curves
shown in this figure are based on measurements on C02 samples
extracted from the air.
The ocean record is based on radiocarbon
measurements made on growth ring dated corals.* The ~ 4 C difference between the atmosphere and surface sea curves provides ameasure of the net input of 14 C to the sea. Part of this input isof course the steady state flux of natural 14 C.
To get the net
flow ~f bomb-produced 14 C, a correction must be made for the airsea b 4 C difference associated with this natural flux.
Of course
in order to calculate a flux, a C02 exchange rate must be assigned
to the particular region of the ocean.
The concentrations of 85 Kr in surface ocean water should beeverywhere very close to equilibrium with the partial pressure ofthis gas in the atmosphere (as is the case for 02). Thus, in constructing models designed to reproduce the observed distributionof 85 Kr in the sea, the assumption is made that the surface ocean
mixed la¥er is saturated with regard to this gas.
The concentration of 5Kr in surface water is then computed from a knowledge ofthe temporal history of the partial pressure of 85 Kr in the atmosphere (and the temperature of the water).
This history (see figure 8-4) is documented by large numbers of radioactivity analyses
on Kr gas separates made in air reduction plants.
The situation
for freon - 11 and freon 12 is much the same as tha t for 8 5Kr .
However, the history of the atmospheric partial pressures of thesetwo compounds is not documented by measurements.
Rather they arereconstructed from a knowledge of the world manufacture of theseproducts.
As freons have no place to go except the atmosphere andas their destruction rate in the atmosphere is very slow, thisprocedure should be valid.
Curves showing the atmospherichistories reconstructed in this way are shown in figure 8-4.

HYDROLOGY OF THE MAIN THERMOCLINE
Before discussing the distributions of these anthropogenic
tracers wi thin the sea, it is necessary to set the stage by summarizing some of the pertinent information about this reservoir
gained by more traditional observations.
Of these, the most*As for trees, annual growth bands can be identified in corals.
This is done by cutting a several-millimeter-thick slab along theaxis of corallite growth through a large coral head and then performing a medical x-rayon the slab.
As the CaC03 precipitated bycorals shows seasonal variations in density, the x-ray absorption
varies producing growth bands on the film.

388

70

NORTH
TEMPERATE

~

_

lJ)

Z

o

U

+ 500

N.TEMP.

- - - EQU.
...... S. TEMP.

3

:E

VI

I-

U

«

v

U

<J

~

I

<I 250~ 250

200-.--.,--...... ~ g U4.
100~ ~
~ <I
l/l

~U

o

01--_":::::"'_--

0 L-.&.-&-~";""a,,,...................I:':~-.L. ..........~..........L.......L.~L.......L...L-L-I-<I
1o....I.-&-"'='55::L--1L.......L.~~60~Y-E ....
A.&.R~.a......L.-&-.l......L~ ...............I....J-100
55

YEAR

60

60

750

65

70

750
SOUTH

EQUATORIALTEMPERATE
~

i!-

500

~

~

:E

500

:E

I-

I-

«

U

U

~

«

<t

<J 250

<1250
o

IOO~

3

l/l

o vU

-<I

'--I..-&-~55~""""''''''''''6i:-::0~-L...l......L6=-=5''''''''-&-~7~0L.......L.--'--L...J-100

YEAR

0l-_ _~ _ - ~<I...................~-L.. ........--":':~--L. .........~........."--L..~L.......L....I-L.J-'OO·
YEAR

Figure 8-3. Plots of fj.14 C for atmospheric CO 2 and for surface
water EC02 from 1952 to 1974.
Separate curves are given for thenorth temperate zone (upper left hand panel), for the southtemperate zone (lower left hand panel) and for the equatorial zone
(lower right hand panel).
The scale for the ocean curve has beenshifted in order to eliminate the air-sea 14 C/ C difference
required to maintain the flux of na tural radiocarbon into thesea.
The differences between the three atmospheric curves stemfrom the finite mixing time of the atmosphere.
As the tests wereconducted mainly in the northern hemisphere, the north temperate
atmosphere shows an earlier and stronger maximum than is seen in
the other zones.
In the upper right panel the difference between
the atmosphere and surface ocean t:::. 14 C values are plotted against
time.
These curves show the extent to which the driving force for
14C entry into the sea changes with time in the three zones.
The
zone to zone differences are quite small.
The equatorial zone has
an integrated dri ving force exactly equal to the average for thenorth and south temperate zones.
The radiocarbon trends for the atmosphere are based mainly on
measurements made by Nydal and his coworkers in Trondheim (7,10,24) and by Munnich and his coworkers in Heidelberg (9,25).
The
surface water curves are based mainly on measurements made during
the GEOSECS program (21,22,23) and on the ring-dated coral measu remen t s by Druf f e 1 ( 15 , 16 , 17).

389

0

E

"[600
"0
rt')

0

->-

~400

-I->

U

«

~ 200

~
LO

co

0

-

NOt

1950

4

1960

1970

1980

YEAR

=>

en
en
w
a:

0

z

0

tU

a<:)

:3

a=>

150 ~

«

I-

~

0

a:

~

a:
a..

LL

w

2000:

2

100

,...,

U

LL

w

>
-t-

50

a::
J:

eL.

(/)

.....J

0

':J

=>
u

-

w

<t
~

E

0'

,...,cn
I
U 0
U
U

U

-

~

0

1950

1960

1970

0
/980

~

YEAR
Figure 8-4. Plots of the 85 Kr and freon concentrations in the atmosphere as a function of time.
The 85Kr curve is based on direct
measurements (107,110).
That for CCl 3F curve is calculated fromdata on the manufacture of this substance.

390important is depth structure of isopycnal horizons.
As we shallsee, this structure provides the basis for subdividing the thermocline into a number of regional uni ts.
It als 0 provides clues tothe manner in which water moves through these units.
The topography of a number of key density horizons along sections through the western sides of the three major oceans is shownin figure 8-5. The most striking feature in these sections is the
upward bulge of these horizons in the equatorial zone.
Anotherprominent feature is the rise of the isopycnal horizons toward thesurface in the latitude range 40° to 50°.
Had these observations
been made during the peak winter months, these horizons would havereached all the way to the surface.
It is these topographic features which provide the basis for our thermocline subdivision.
Ineach ocean we designate equatorial and temperate zones.
Theboundary between these zones is taken to be 15° (the latitude atwhich the slope of the isopycnal surfaces is the greatest).
The
poleward boundary of the temperate zones is taken to be 45°.
At
this latitude the key horizons are rising most steeply.*
Another important source of information comes from B-3 diagrams.
Such diagrams are shown in figure 8-6 for key s ta tions
along the above sections.
To understand the point of these diagrams one aspect of the ventilation of the main thermocline must
be mentioned.
It is the opinion of mos t physical oceanographersthat at least in the temperate ocean, water reaches the interiorof the thermocline along isopycnal horizons (i.e., new thermocline
water is generated during winter months when the isopycnal horizons reach all the way to the ocean surface).
A map showing thelocation of these outcrops in the North Atlantic is shown in figure 8-7.
This gives rise to an interesting situation:
there are two
sets of outcrops for each isopycnal, one in the northern ocean and
one in the southern ocean. Where then, along each isopycnal horizon, does -the boundary between the two water types lie? Or is
there a continuous gradation in composition along each surface?
As seen in figure 8-6, the B-S combination producing water ofa given density is quite different in the north than in the south.
In the Atlantic the northern waters are more salty than the southern waters.
Hence the temperature of the water on any given isopycnal surface is warmer in the north than in the south.
A map oftemperature along the 26.75 0 /00 isopycnal surface in the Atlantic is shown in figure 8-8.
A sharp boundary between waters ofnorthern and southern origin is seen in the vicinity of 15°N.
In
this ocean ~t least the thermocline of the equatorial zone is supplied primarily by waters originating in the south!
In the Pacific the situation is the reverse of that in the
Atlantic.
The northern waters are considerably less salty than
the southern.
As in the Atlantic, the equatorial waters are morenearly akin in their 8-8 characteristic to the southern than to*I'he Indian Ocean does not have a full-fledged, north temperatezone. Because the Indian Ocean is bounded at ~ 20 N by the Asian
continent, its isopycnals do not reach the surface even during thewintertime. This leads to the very low 02 concentrations seen inf i gu r e 3- 11.
0391

WESTERNATLANTICDENSITY

~

I-

a..
w
o

1.5 "-~~_&.-_""l--I~.....:II~-.a--l."""'_"""'_"""'''''''''''
40
20
a
20
LATITUDE
N
0.0 ~'P"""""'....--...--....-'P"""""'_"""''''''''''''''~_

s

:J:
I-

a..
w
o 1.0

40

20

0

LATITUDE

20

N

Z
I-

a..
w
a

s

o
LATI TUDE

20

N

Figure 8-5. Sec tions along the western side of the world's threemajor oceans showing how the depth of various isopycnal horizonschanges wi th la ti tude.
The heavy dashed line shows the depth atwhich the tritium concentration of the water (at the time of the
GEOSECS surveys) was one quarter that of the overlying surface
water.
Based on data obtained as part of the GEOSCES program(424,425,426).

392

20,...-......- -......-......--..---...,.-,.,....,

WESTERN
ATLANTIC

u
o
0:

0:

TREND\ix(ALSO
EQU.
TREND).-::><t:
W

a..
~
w

o

W
0:

SOUTH
TEMP.

W

Ul5

::>

~

0:
wlO

a..
:E
w

.-

.-

...J

...J
<t:

_ 27°NZ

o 24°N.-<t

W* 4°N-4°S

O

El 24°S
.27°S

t-

a..
34

35

..-o
Z
W

a..

36

0-------.. . - ....--.. . - ....---..
34

37

SAL IN ITY(%o)

twl5

WESTERN
PAC I F IC

0:

::>

....
<t

~IO
~
w

t...J

<t
Z

w

NORT H
TEMP.
TREND

/4
~

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0:14

*r.1~ TEMP.

5.-<t:::>

TREND

a::

Wl2

0°
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a..

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EQU.
TREND

w

t..JIO .

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El 24°5• 27°So-----"-----oA--.. . . ---'
33

34

35
SALIN ITY (%0)

W. PAC.

t-* 4°N-4°5

a..

'.

;,

W

o 25°N

o

,

<{

....Z

_ 30 0 N

....

36

~O

fa'} ~fa·
1,'-*Jt:(-SOUTH.~ 7.-~.

a::

....

I-

':>

35

SAL INITY(%o)

36

45

S

30

15

0

,

~

15

LAT ITUDE' .....

----.,.,
30

45

N

Figure 8-6. 8-S diagrams for the thermocline waters of the threemajor oceans.
In each case the trends for north temperate,equatorial and south temperate waters are shown separately.
Alsoshown are several key isopycnals.
In the lower right hand panel
the temperature of water along the 26.75 % 0 isopycnal horizon
is shown as a function of latitude in the western part of eachocean.
Based on data obtained as part of the GEOSECS program(424,425,426) .

393

60

40
w

o

:::>

t-

~ 20

«

..J

o

20
100

80

60

40

20

o

20

LONG I TUDE
Figure 8-7. Map showing the winter outcrops of isopycnal surfaces
in the North Atlantic Ocean.
This map was cons tructed from the
NOAA Geophysical Fluid Dynamics Group's oceanographic data files(689) by Jorge Sarmiento of Princeton University who kindly
supplied this diagram for use in the book.

394

the northern water.
In this connection it should be noted that in
the deeper parts of the equatorial thermocline, the water on any
given isopycnal has a higher salt content than either the water tothe north or the south. Thus they could not have been produced bymixing of waters generated at the outcrops of the isopycnal surfaces.
One explanation for this seeming dilemma is that the 8-S
structure in the equatorial zone has been modified by verticalmixing (at least in density range 26.7 to 27.7 0 /00).
The situation in the western Indian Ocean is similar to thatin the Atlantic.
The northern waters are saltier than the southern waters and the upper thermocline water in the equatorial zone
are more akin in their 8-S relationship to the southern than to
the northern waters.
In the lower part of the equatorial thermocline, however, the composition moves away from the southern trend
and becomes more similar to the northern trend.
The high salini ty of waters in the northern Atlantic andnorthern Indian Oceans must have their origin in large net evaporative losses from these basins.
We emphasize net because, werethe evaporated water to rain back onto the basin or onto continental areas draining back into this basin, the salinity would not beenriched.
There must be a loss of vapor which ends up falling as
rain onto the North Pacific Ocean.* This is especially surprisingin light of the fact that 60% of the world's river runoff reaches
the North Atlantic!
Somehow an even greater loss of water vapor
from the Atlantic across Central America with the trade winds andfrom the Mediterranean across the Middle East to the Indian Ocean
and its drainage basins must occur.
If the route from theMedi terranean to the Indian is important, then a corresponding
loss must occur from the northwestern Indian Ocean carrying thevapor to the North Pacific Ocean.
By contrast the salini ties on
any given density horizon for the three south temperate regions
are quite similar.
This must be a consequence of interoceantransport via the circumpolar current.
In any case, these idiosyncrasies of the global water vapor cycle provide the oceanographer with an excellent tracer for thermocline processes!
Another set of observations having importance with regard to
the thermocline is the pattern of the overlying surface currents.
These currents often have roots carrying thermocline water along
the same paths.
Due to the rotation of the earth and the patternof the winds, the temperate zones are characterized by strong
western boundary currents which carry water poleward.
These intense currents bend around to the east and carry waters across theocean along the poleward boundary of the temperate zones.
The
Gulf Stream in the North Atlantic and Kuroshio in the North Pacific are currents of this type.
This circulation gyre is closed by*We exc lude transport to the sou thern ocean because the res idence
time of water vapor in the atmosphere (~ 2 weeks) is small comparedto the interhemisphere mixing time (~ 1 year), and because the salt
deficiency in the North Pacific nearly matches the salt enrichment in
the North Atlantic and North Indian Oceans.
The southern oceans are
similar to one another in their 8-S relationships and lie in between
the North Pac ific, on one hand, and the North At lant ic and North
Indian, on the other.

395

o

c:,A,)c.o0')Figure 8-8. Maps showing the distributions of potential temperature and of tritium
concentration along the 26.80 % 0 isopycnal horizon in the Atlantic Ocean.
The
shaded area shows the region of transition between waters with southern and northerncharacteris tics.
Based on data obtained as part of the GEOSECS program by Gate
Ostlund of the University of Miami (424).

a more dispersed equatorward flow through the central and easternparts of the zone and finally by awes tward flow through the low
latitude portion of the temperate zone.
These same currents are
known to operate at greater depths.
Thus they must carry the newwater entering the thermocline each winter around a similar pathway.
From the sharp change in potential temperature along any
given thermocline isopycnal seen at 15°N in the Atlantic, it canbe inferred that this lateral stirring is confined to what we havedefined as the temperate zone.
Within the equatorial zone there are strong currents runningback and forth parallel to the equator.
Some of these currents
are strongest at the surface and some are strongest beneath thesurface in the upper thermocline.
While we need not be concerned
wi th the details of these currents, we must be aware of the factthat they very effectively carry water (and our tracers) acrossthe equatorial zone and back on the time scale of one year.
More subtle are the currents which carry water from one zone
to another, and even more subtle still are those that carry waterup or down (i.e., across isopycnals).
One might say that they arelost in the "noise" of strong lateral currents which recirculate
water within a given zone.
We do, however, have some clues to theexistence of these currents.
One such clue comes from the geographic pattern of potential
temperatures along isopycnal horizons in the thermocline of theNorth Atlantic.
It can be seen in figure 8-8 that the waters in
the 12° to 16°N zone have temperature transi tional between thosefor northern and southern water.
Thus it must be in this zonethat the two water types mix.
This transition zone extends
through the Caribbean, the Gulf of Mexico and the Florida Straits
into the temperate Atlantic (see figure 8-8). This pattern gives
the impression the water feeding the roots of the Gulf Stream consists in part of recirculated north temperate water and in part ofwater moving up the Atlantic from the south.
As such interzonal
flow must occur in order to balance the export of the deep waterproduced in the northern Atlantic, it is not surprising to see
evidence for the intrusion of southern water in the upper ocean.
Evidence for upwelling is provided by several chemical signatures.
We have already discussed the excess partial pressures of
C02 and N20 gas and the high nutrient constituent concentrations
in surface waters of the equatorial Pacific.
These anomalies canbe maintained only by upwelling.
The high opal content of sediments beneath this zone also provide evidence for upwelling (seefigure 2-2) ~Another upwelling signature is provided by higher
than average rates of plant activity. A section of plant product\vity (as measured by the rate at which plants take up artifical4 C) across the equatorial belt in the Pacific Ocean confirms what
we have already inferred from the chemi cal resul ts (s ee figu ree-9).
The high nutrient constituent content of this water supports the high plant growth rates.
The map in figure 8-10 shows a global compilation of plant
growth rates.
Zones of high productivity are shown by the crosshatching.
These zones are found along most continental margins
and in most high latitude regions, as well as along the equator in
the Pacific.
The zones of high productivity in the polar regions

397

"""'"'

0

-E 50
J:

.....

EQUATORIAL PACIFIC
112 0 W

a..

w
0

100

5

a

5

5

LATI TUDE

N

Figure 8-9- Plant producti vi ty (in uni ts of 10- 3 gelm 3 day ) as afunction of latitude and depth along a section through theequatorial Pacific Ocean.
Adapted from a section in reference687.

< 4.5 mol/m 2 y.,

4.5-7.5 mol/m 2 yr>. 7.5 mol/m 2 r

Figure 8-10. Geographic distribution of plant productivity in theocean.
Adapted from a map in reference 686.

398

owe their orlgln to the outcropping of isopycnal horizons.
Nutrients are returned to the surface along these horizons of preferred mixing.
The zones of high productivi ty in coastal areas
owe their origin primarily to wind-driven upwelling, although thedelivery of new nutrients from rivers and the efficient local
recycling of nutrients from shallow water sediments also contribute.
Our interest is of course in that part of the high productivity supported by upwelling.
Physical oceanographic studies
have clearly demonstrated that upwelling is the major cause forthe high productivities in the eastern Atlantic, eastern Pacificand western Indian Oceans.
The warm surface water capping thethermocline in these regions is periodically blown seaward drawing
the underlying cold waters to the surface. This wind-driven coastal upwelling is one of the major routes by which thermocline waterreaches the surface at low latitudes. Before completing the discussion of the map in figure 8-10, the reader should no~e that thelower limit in the high-productivity zones (7.5 moles/m yr) is noteven a factor of two greater than the upper limit for the lowproductivity zones (4.5 moles/m 2 yr).
As differences in the techniques used to measure productivity may lead to systematic errorsof 15% or more, the distribution of productivity is not nearly so
precise as that for chemical properties which have been measured
ocean wide by a standard technique.
Thus the map is bes t viewed
as a broad brush view of the contrasts in plant productivity
across the world ocean.
While the major features are surely correct, the details are suspect.

TRITIUM DISTRIBUTION WITHIN THE THERMOCLINEWith this background of thermocline structure and dynamics in
mind, let us have a look at where the anthropogenic tracers )lavegone.
To date we have only global coverage for 3H and for
4C•
As we shall see, the limited data on 90 Sr provide a valuable supplement to the more extensive 3H data set. Surveys of the distribution of 85 Kr and of the freons have only recently begun.
Another decade will pass before global data sets are available forthese gases.
We begin this discussion with the distribution of tritium.
The reason is that, as the contribution of prenuclear tritium isvery small, the anthropogenic patterns are immediately apparent.
Sections identical to those presented for densi ty in figure 8-5
are presented for the distribution of tritium in figure 8-11.
Astransport "along isopycnal horizons is thought to be very important, this figure should be compared wi th figure e-5.
To aid inthis comparison, the dashed line showing the depth at which thetritium concentration reaches one-quarter of the value for overlying surface water is shown in both sets of diagrams.
The most dramatic aspect of these diagrams is the asymmetrical nature of the distribution of this isotope about the equator.
The north temperate oceans have far more tritium at any density
level than do the south temperate zones.
This is of course to be
expected from the input function for tritium.
The integrated
water inventories of tri tium at the time of the GEOSECS surveys
dramatically portray this asymmetry (see figure 8-12).

399

E
.¥

-

::t:

I-

WESTERN
ATLANTIC
1972
TRITIUMtu0Wo1.5'-~~-""'~~""'.a-_*0--I.......L.~~-I....I.J.~=:--...J

LATITUDEN

E 0.5..:.:::

~

I0W

o

J:

I-

WESTERN
PACIFIC

L1J

1973-1974a.
C

TRITIUM

tu

1.5

0

S

LATITUDE

N

Figure 8-11. Sections showing the vertical distribution of tritium(in T.U.) along traverses in the western part of the three major
oceans.
The heavy dashed line denotes the depth at which thet~itium
concentration~eaches
one-quarter thatin overlying
surface water. These results were obtained by Gate Ostlund of the
University of Miami as part of the GEOSECS program (51,52,54,56).

400

S

LATITUDE

N
40

a

40
l:l EASTo WEST

40

20

10

>a::

r:P e

a

~

z_

WC\J

>E

0
8
INDIAN

~z~
~

(/)

-ZE
~~o

0:::)0J- ....J cn

°0
ua::
w

t<{

o

OCEAN

0

1977-1978

0

:~

1

s:>

0

~ ~dt"000..
~

I . " _

40~

PACIFIC

OCEAN

0

~

1973-1974

30

0

e
0

00

#0

a

0

o

19 ~
'----.l.....-.---'~,I,;",.,I,....;;.--I._........L_....L._...J.._...

40

20

S

40
N

Figure 8-12. Water column inventories of tritium as a function oflati tude in the three major oceans at the time of the GEOSECS
surveys (51,52,54,56).

401·
A second prominent feature is the deeper penetration of tritlum in the temperate zones than in the equatorial zones.
As thisfeature is obscured by the asymmetry of the distribution it isinstructive to observe the shape of the horizon at which the tritium content reaches 25% of its value in overlying surface water.
Except for the northern Atlantic, the penetration in polar
regions is not as grea t as in the adjacent temperate regions.
Inthe northern Atlantic the tritium is found all the way to the bottom.
This penetration is of course related to the formation ofdeep water in this region. The lack of penetration in the Antarctic and northern Pacific is related to the capping of theseregions by low-salinity waters.
Only in special places (like theWeddell Sea) is this stratification broken, permitting deep penetration of tritium.
A summary of the average penetration depth* of tritium forthe central parts of each zone (i.e., away from the zone oftransition separating these thermocline units) is given in table
8-2.
The penetration is greatest in the north temperate zones,
intermediate in the south temperate zones and smallest in theequatorial zones.

TEMPORAL TRENDS IN TRITIUM
Beyond knowing how the distribution of tritium looked at one
single time, it is important to know how it evolved through timeat any given place in the ocean. Two aspects are of interest, themanner in which the mean penetration depth changed and the mannerin which the water colunm inventory changed.
At the time thisbook was written there were only two places in the ocean where we
had adequate data for this purpose. **One is in the vicinity of
the island of Bermuda in the north temperate zone of the Atlantic.
As shown in figure 8-13, the decay-corrected tritium concentrationin the Bermuda area shows remarkably li ttle change between 1968
and 1977.
The concentration is uniform to a depth of about 400
meters. Between 400 and 800 meters there is a linear drop-off inconcentration.
From 800 meters to 1200 meters the values areagain roughly uniform.
A closer examination of these data reveals, however, that the concentra tions in the upper 400 meters
have declined, while those below 400 meters have increased.
In
terms of the water column inventory the two changes roughly balance.
The small increase that is suggested by the resul ts (seetable e-3) could in part be explained by inputs during thisperiod..
(about 6% of the total input).
Thus the north temperatezone did not lose an appreciable amount of its tritium throughmixing with low-tritium southern waters between 1968 and 1977.
*The average penet ation depth is defined as the2inventory (atoms/em) divided by the surface water3(a toms/ cm ) ..

water column
concentration**Short ly after the pub 1 icat ion of this book the resul ts from thenorth and equatorial surveys from the Trans it Tracers program will
become available, greatly increasing our knowledge of the evolution
in time of the tritium distribution.

402

Table 8-2. Mean penetration
the GEOSECS surveys.*

Ocean(Year)

depth

N. Temp.
35°N
to

of

tritium asobserved during

Equat.

S. Temp.

lOoN

25°S

to

25°N

to

laoS

35°S

Bomb Tritium
Mean Penetration DepthmetersAtlantic (1973)
Indian(1978)
Pacific(1974)-

680 (5)565 (11 )

135 (8 )
180 (12)195 (18 )

345 (3)390 ( 4)330 (7)*The mean penetration depth is obtained by dividing the water c~l­
umn inventory (atoms/m 2 ) by the surface concentration (atoms/m ).
The numbers in parentheses are the numbers of stations included toobtain the ave rage.
The tri tiuffi data was obtained by Os tlund of
the University of Miami (51,52,54,56).
The compilations summarized here were made by the authors (75).

403

BERMUDA AREANORTH ATLANTIC

*

(::J

(::J~
~

I>

+c:J

0

/A x ~0

1968

Y

:c

I

/~

~:

x

..
S

~

a..

lJ.J

o

.. 9/68
x 5/72+ 9/72
I:l

x0E1

~

3/77

o 5/77~

6/77* DECAY CORRECTED
TO 1/1/1973

.1.2

x

+

<8)

8-13. Tri tium profiles in the Bermuda area in 1968, in
1972, and in 1977. The concentrations are decay corrected to thebeginning of 1973.
The 1968 and 1972 results were obtained by
Gate Ostlund of the University of Miami (49,51) and the 1977

FigureresultsbyBill
Institution (55).

Jenkins

of

the

404

Woods

Hole

Oceanographic

00....0_ _0....2_.---0...4_ _0.......
6 __0.....;.-.8_ 8 1.0·

CALIFORNIA CURRENT
28°N 122°W

.'18
4*,
;

.8

•

I

t- .2

3

0W

£:)

• SEPT. 1969

* NOV. 1971

A JUNE 1974

1.2*

H SURF H20
(t u)

23.0
17.4

12.2

.4 - - . - -.....- -......- -.....--~--.....- -....
Figure 8-14. Profiles of tritium taken in 1969, 1971 and 1974 at a
station in the California Current (48,54,58).
The values are
given as the fraction of the average value for surface mixed layerwater.

405

The second locality is in the California Current adjacent to
the Mexico- Uni ted Sta tes border.
This current flows down the
coast carrying low-salini ty water similar to that found off the
Aleutians. The situation at this locality is quite different; thedecay-corrected water column inventory shows a dramatic decreaseover the five-year period (1969 to 1974) and the mean penetration
depth remains almost the same. The tritium profiles for this station are shown in figure 8-14.
The concentrations are expressed
as ratios to the mean for samples from the upper mixed layer.
This mode of presentation facilitates comparison of the shapes ofthe depth profiles.
The mean penetration depths and decay-corrected inventories are summarized in table 8-3.
It appears thatin this area of the ocean, tritium was being lost from all depths
in about the same proportion.
As part of the water carried in thesouth-flowing California Current may be transported into the equatorial zone, some of the missing tritium may have been lost entirely by the temperate zone.
The remainder is carried around thelower part of the gyre into the western part of the temperatezone.
For some other parts of the ocean enough data from different
times is available to permit rough estimates of time trends in inventory and mean penetration depth to be made.
These are summarized in table 8-4.

TRITIUM AS A GUIDE TO DEEP WATER FORMATION
Before leaving the discussion of tritium, we must briefly
mention its application to the study of deep-water formation.
Asshown in figure 8-15, tritium has penetrated all the way to thebottom of the northern Atlantic.
Particularly well shown is thecascade of new Denmark Straits overflow water along the bottom
down into the abyssal western basin.
A program was in progress at
the time this book was wri t ten to repeat the survey along thistrack nine years later (in 1981) to determine the extent to which
this penetration has progressed.
The existence of this thin river of bottom water is very
nicely shown by the profiles of temperature, salinity, lightscattering, dissolved oxygen and dissolved silicate at a station
located at 39°N in the western basin (see figure 8-16).
Here a25-meter-thick layer of water which is more salty and slightly
warmer than the ambient bottom water is seen.
The water in thislayer also contains more dissolved oxygen and less silicate thanthe overlying water.
Its higher tritium content demonstrates that
it is richer in new northern component water than the overlying
water.
There is evidence that this stream of deep water flows wellbeyond this la ti tude.
It is thought to hug the margin of theNorth American continent as far south as Florida!
As shown infigure 8-17, Bill Jenkins and Peter Rhines of Woods Hole have provided dramatic evidence for this continuation based on tri tiumdata.
At about 30 0 N this deep south-flowing stream is thought toleave the continental rise.
As it does so, it loses its sedimentload (torn loose from the slopes it has hugged).
This sedimenthas built a huge abyssal delta called the Blake-Bahama Outer Rise.

406

Table 8-3. Time trends in the water column inventories and mean
penetration depths for bomb produced tritium in various places in
the Atlantic Ocean.

Stn.

Col. Date

Lat.

Long.

h

m

WCI*
10 9 atoms
cm 2

Ref.

550
600
750
750
700
740
770
770

28
28
33
34
34
34
33
32

( 49 )
(49)
(51)
(55 )
(55)
(55)
(55)
(55)

45
60
105
120

1.1
1.3
1.3
1.7

(47)
(47)
(51)
(51)

210
210
210

29
24
19

(58)
(48)
(54)

TEMPERATE NORTH ATLANTIC
P6808-2
BRMD-72
GEOSECS 30

S74546
S75625

PAN001
PAN002
PAN003

Sept 1968
May 1972
Sept 1972
Nov 1974
Mar 1975
Mar 1977
May 1977
June 1977

31.0 0 N
32.3°N
31.8°N
30.5°N
30.5°N
30.2°N
30.2°N
30.2°N

64.9°W
65.5°W
50.8°W
67.5°W
58.5°W
64.5°W
64.5°W
64.5°WEQUATORIAL ATLANTIC
P6706-48
P6707B -12
NAGS-8NAGS-3B

Sept 1967
Nov 1967
Sept 1973
Sept 1973

0.0°
0.0°
5.0 0 S0.2°S

4.0 0 W
18.2°W
11.1°W
10.0 0 W

CALIFORNIA CURRENTGEOSECS IGOGOGEOSECS 347

Sept 1969
Nov 1971
June 1974

28.0 0 N
28.0 0 N
28.0 0 N

121.0 0 W
121.0 0 W
121.0 0 W

EQUATORIAL PACIFICSTYX 3
GEOSECS 239

Aug 1968
Dec 1973

7.0 0 N
5.9°N

163.0 0 W
172°0 0 W

+
+

6.3
6.8

(50)
(54 )

STYX 2
GEOSECS

Aug 1968
Dec 1973

0.0°
0.0°

165.7°W
179.0 0 W

200
200

6.3
4.0

(50)
(54)

STYX 1GEOSECS 251

Aug 1968
Dec 1973

5.0 0 S
4.6°s

167.1 0 W
179.0 0 W

160
210

2.4
4.0

(50)
(54)

*The water column inventories have been decay-corrected to Jan. 1
1973.
+Profile has subsurface maximum.

407

Table 8-4. Summary of information regarding the time trends in the distributions ofbomb-produced tritium and radiocarbon at various places in the ocean after 1967.
The
dashes indicate a lack of information.

TRITIUM(Decay corrected)
Concentration
Mean
Surface
Penetration
Ocean
Depth

RADIOCARBON
Water
ColumnInventory

Concentration
Surface
Ocean

Water
Mean
Penetration Column
Depth
Inventory

NORTH TEMPERATE ZONE
Atl.
Ind.
Pac.

-3%/yr-7%/yr+4%/yr--+::ao-Constant-Constant-Constant-Constant

-

-

-

-

-

EQUATORIAL ZONE

o

00

Atl.
Ind.
w. Pac.

-8%/yr+3%/yrIncrease
Increase

Increase
Increase-Constant-Constant

-

-

-

-

-

-

-

SOUTH TEMPERATE ZONE
Atl.
Ind.
Pac.

-

-

-

-

-7%/yr

-

-

-

-

-

-

BOUNDARY CURRENTS
California
Aleutian-t5%/yr-ll%/yr-Constant
Increase-9%/yr-7%/yr-Constant

-

-

1.0

2.0

----E
~
.,.,....
J:

I- 3.0

a..
w
0

EO

lOON 20° 30° 40° 50° 60°
LATITUDE

Figure 8-15- Section showing the vertical distribution of tritiumin the western North Atlantic.
The deep penetration north of 40 0 N
is associated with the formation of deep water.
This diagram was
prepared by Gate Ostlund of the University of Miami based on his
measurements on samples collected as part of the GEOSECS program( 727) .

409

4.5 r----"r----.-..-"r------.,r-----r--,_

. :1'
.,. "
4.61-

.'

,:

~,

....
. ....

-""

E
oX 4.7 l--."~::::'"

I

l-

n..

"

-

~ 4.7 ~

-

I

I-

.

~ 4.8 ~~.~~ ~

a..

J

". :

~ 4.8-

4.6~

-

-

"

'.'
.: .

~..

• I .'"-NEPHELS

-

....
4.,
~

"

4.9 -

#

,y

IC.·-

#

-

4.9 ITEMPERATURE

•

5,0

•

I

34.89034.895I

34.900

SALtNITY(%o)

•

I

5.0
2.250

34.905

2.275

2.325

2.300
TEMPERATURE (oC)O2 (lLmol/kg)

1
0

I

32

I

J:

a..

I

l-

0

w 4.80

/

0

0

/

4.85
0"

1.79

/.81

\

I

H4 Si0 4

1.82

0

~

0

en
~

\

w

c/O

0

Z

/

U

0/

:2

0
f- 150 f-

\

<t

~/d

\

~ 4.80-

[]J

I

\00 \

I-

\

~
(f)

4.851-

02

0,

s

w

>

- 1000[]J«\. . aD

w

u

\

0

-~

(5

a

__

34.89

- 50 Z«I-(f)o-:::.-=--o~

I

I

1.83
POTENTIAL TEMPERATURE (OC)

200'E

\

\

a..

,

\0

I

E 4.75 I-0\\

oX

6

275

\

\

w

fOO

,

.p \ 0 •

~

en

/0
/
0
...-IT_ ~ - - ~ 0-0, - -

1.80

4.70

0
150 ll0

/0

')01

200

E

/
POT. TEMP.

270

I

"'\ \
00

a

a

34.90

SALINITY("!oo)Figure 8-16. Profiles of salinity, temperature and light scattering for a station in the northern part of the Atlantic's westernbasin (39°N, 44°W) and obtained using in situ sensors (upper
panels) .
Also shown are conventional, dis crete measurements ofpotential temperature, silicate, dissolved oxygen and salinity(lower panels) (429).
In all cases an anomaly with characteristics more like that of water entering the northern Atlantic fromthe :Denmark Strai ts is seen as a several tens of meters thickbottom layer.
This layer carries warmer and more saline waterwith higher dissolved oxygen and lower silicate than the overlying
water.
It also scatters less light (i.e., presumably has a lower
particulate load).
Its tritium concentration (not shown) is
higher than that for the overlying water.
These results wereobtained as part of the GEOSECS program (362).

410

- - - - - - - - - - 8=3.0----.,........

E~......,
I

----

~

a..

w
0

4
~.)

40·

Figure 8-17. Tritium section for bottom waters flowing along theBlake-Bahama Outer Rise (location B inset map). . The solid
contours are for tritium (in T.U.) and the dashed contours are forpotential temperature.
The location of the station from which therecords shown in figure 8-16 were obtained is designated in theinset by point A.
The tritium measurements were carried out by
Bill Jenkins of the Woods Hole Oceanographic Institution (70).

411

As shown in figure 8-17, a plume of tri tiurn-bearing water flows
along this feature.
Al though the concentrations are lower, thistritium anomaly remains identifiable 3000 miles from its source!

SUPPLEMENTARY INFORMATION FROM Sr-90
We have already pointed out that the distribution of 90 Sr in
the sea is quite similar to that of 3H.
Because the tritium database is more complete, and because the precision and sensitivityof 3 H measurements are superior to those of 90 Sr , it is the tritium data base which receives most attention.
There is one aspectof the data set for 90Sr, however, which is superior to that for
3H .
The time histories for the concentrations of 90 Sr in surface
water in the latitude range about 20 0 N to 20 0 S can be reconstructed in far more detail than those for tritium. The rea~on isthat the concentration of 90 Sr is recorded by corals.
The 9 Sr/Sr
ratio in corals must be the same as that in the water in which thecorals grew. *The corals neatly average out any seasonalvariability so measurements on individual annual bands remove someof the noise that would be seen if, instead, spot water samples
had been analyzed.
More important, as the record from spot water
samples is generally inadequate due to a sparsity of coverage, it
is possible to make up for this missed opportuni ty to collectwater samples by collecting coral heads from sites of particularinterest.
This work had only just begun as we wrote this book, sothat its full potential has yet to be realized.
The four coralbased records obtained by Toggweiler at the Lamont-Doherty Geological Observatory and shown in figure 8-18 demonstrate how powerful
an approach this is.

BOMB C-14 DISTRIBUTION WITHIN THE THERMOCLINE
Bomb-produced radiocarbon cannot compete with the tritium asa tracer of the details of the penetration of anthropogenic substances into the sea.
The reason is that the task of separating
the contributions of bomb- and cosmic-ray-produced radiocarbon
cannot be done well enough.
The value of bomb radiocarbon comes
rather through an unexpected contrast in its water column inventory between the equatorial and temperate zones of the ocean.
In order to obtain these water column inventories, we must ofcourse make the separation of the two types of radiocarbon.
Themeasurements made as part of the GEOSECS program provide the profile of total (i.e., natural + bomb) 14 C to C ratio at many placesin the ocean.
The trick is to es tablish the prenuclear profile
for these same places.
As we have only fragmentary data, we cannot simply plot measurements for samples taken at various depths
and draw a smooth curve through them.
A less direct method must
be employed.
The procedure used is as follows.
First the surface
value is established from measurements on prenuclear water samples
or on growth rings from corals.
Next the GEOSECB tritium data are*As the. Sr / Ca :- a t i ~ i ~ cor a l s i s near 1y the s arne a s t9~a tin ~ e awater, In practlce It IS necessary to measure only theSr actlvity per gram of coral.

412

70

70

C1a

PACIFICOCEAN

ATLANTIC
OCEAN

60
50

0,

0
0

50
OAHU 21°N

00...... 40
E

, 40

.~

E

0-

0"0

~

-:: 30

'U>

U>

0

0

0

(1)

(1)

O----..a----.--.....-.-------.a..--~

1960

1970

1980

O'---------------..a.--------.--..-~
1960
1970
1980

YEAR

YEAR

Figure 8-18. Records of 90 Sr concentration as a function of timein surface waters as reconstructed through measurements on growthring-dated corals.
These resul ts l,vere obtained by Toggweiler ofthe Lamont-Doherty Geological Observatory as part of his PhD
thesis research.

413

used to establish the depth of penetration of the anthropogenic
tracers.
In waters which have less than 5% the amount of tritium
found in the overlying surface water, the radiocarbon is assumed
to be entirely of cosmic-ray origin.
Hence, the prenuclear profile must join the profile measured during the GEOSECS program at
this depth.
The final step is to decide on the shape for the
curve joining the surface value with the profile beneath the depthof tritium penetration.
As different procedures are adopted tosuit the idiosyncras ies of various ocean regions, it is bes t tohold off on this point until we discuss each region separately. In
any case, having estimated the prenuclear profile, we use the area
between the total and prenuclear curves to estimate the number ofexcess 14 C atoms in the water column at each of the GEOSECS
stations.
To do this properly we must of course know the amountof dissolved inorganic carbon in the water.
This vari~s slightly
from region to region and from depth to depth.
However, as LCO 2
data are available for all the GEOSECS stations, it is Just a matter of ari thmetic to convert the /),14 C values back to 1 C/C ratios
and to multiply these ratios by the appr0r,riate LC02 values andthen to sum over depth to obtain the bomb- 1 C inventories.
The critical step in this analysis is the establishment ofprenuclear 11 14 C value for surface waters.
Fortunately a reasonable number of water samples collected prior to 1958 were analyzed
for radiocarbon.
These analyses are plotted as a function of latitude for each of the oceans in figure 8-19.
The GEOSEGS data forsurface ocean water are also shown on these diagrams.
The prenuclear analyses appear to follow the pattern seen in the 1970's,
i.e., the highest 11 14 C values are found in the temperate zone,
with a small decline in the equatorial zone and a large decline athigh lati tudes.
The similari ty between the shapes of the two
curves is to be expected.
The processes which reduce the surface
concentration of excess 14 C must also reduce the concentration ofnatural radiocarbon.
In this regard it is important to assess the validity of theearly radiocarbon measurements.
In the twenty years which elapsed
between these early surveys and the GEOSECS survey, many improvements were made in the manner in which the large water samples
needed for this work are collected, stored and stripped of theirdissolved inorganic carbon, and in the procedures used to measure
the radiocarbon in this C02.
Two methods are available for checking the reliability of the early work.
One is to compare the results obtained on deep waters still free of tritium at the time of
the GEOSECS surveys.
Figure 8-20 shows such a comparison between
the early Lamont-Doherty Geological Observatory resul ts on deep
waters in the Atlantic and results for the same geographic location and depth interpolated from the network of GEOSECS measurements.
Considering the experimental error on the early results (~10 0 /00) and the uncertainty in the result interpolated from the
GEOSECS data set (-5 0 /00) the degree of agreement is satisfactory.
The other approach· is to make use of new measurements on
ring-dated corals and on shells from museum archives to determine
the 14 C/ C ratio in surface waters at specific times in the past.
The work on corals looks particularly promising in this regard.
At the time this book was written only three such records had been

414

200

ATLANTIC
OCEAN

0

/0

0
..........0 ...0

4

Q

0

0~

u
:!

<l

0

0

-100

80

N

100

o !

u

<l

-100

_ - . , . _ -... 200

100o

LATITUDE

60

N

Figure 8-19. /:114 C for samples of surface wa ter as a func tion oflatitude in the three major oceans.
The solid circles are formeasurements made on water samples collected prior to 1958 (i.e.,
before the onset of contamination with bomb- 14 C).
These measurements are given in references 4 5,6 and 8. The prenuclear results
obtained by the analyses of coral are shown by xTs (15,16).
The
open circles are the results from the GEOSECS survey (21,22,23).
Because of the sparsity of pre-1958 data in the Indian Ocean,measurements on samples taken in 1960 and 1961 are also shown(open triangles) (13).
j

415

-75

ATLANTIC> 1.4 km DEPTH

• •
(f)

u
w -tOO

•

(f)

0

•

W
<.9

•

...-...
0

~
0

~

U

v

-125

<l

-150

-75

•

EXP. ERRORS

-100

-125

-150

.6 14 C(%0) LAMONT
Figure 8-20. Plot of /114 C values inte~polated from the GEOSECS
results (21) against /114 C values measured two decades earlier onsamples from greater than 1.4 km depth in the Atlantic Ocean (4).
The median line is that expected if the two data sets agree.
The
flanking lines represent the one standard deviation error limits.
The agreement between these data sets bears witness to the reliability of the early measurements.

416

obtai ne d by Druf f e I at Uni ve r sit y of Ca I i f 0 rni a at San Di ego.
They are shown in figure 8-21.
For the two Atlantic corals thepre-1958 /).14 C values are in excellent agreement wi th the earlymeasurements on surface waters.
The coral measurements also
confirm that 1957 is the cutoff date for the acceptance of surface
water results as free of contamination with bomb-produced 14 C•
So we see that our knowledge of the prenuclear 14 C/ C ratios
in surface ocean water rests on a reasonably firm foundatio~ andthat this foundation can be strengthened by measurements of
4 C/ C
ratios on growth-ring-dated corals and on shells carefully selected from museum collections.
Having es tablished the prenuclear surface 14 C/ C ratios fo r
surface water and the depth in the water column at which the GEOSECS results can be taken to represent prenuclear conditions, only
the task of constructing a curve joining these r-esul ts remains.
Here we are guided by our knowledge of the vertical distributionof total radiocarbon in the upper ocean and also by some physicaloceanographic horse sense.
In the equatorial zones of the oceans we have already seenthat the penetr-ation depth of tritium is much smaller than elsewhere in the oceans.
As shown in figure 8-22, plots of tri tium
versus potential temperature generate a linear frray of points.
Although the scatter is greater, so also do the /). 4 C values.
Ourintuition would suggest that during pre-nuclear time the trend of/).14 C through this zone should also have been linear.
We can fortify our intuition by noting that if upwelling is important in theequator-ial zone, then the trends of all conservative properties
being mixed down into the wa ter column against this upward flow
should form straight lines when plotted against potential temperature.
Thus the assumption that the trend of natural radiocarbon
was linear with potential temperature (at least in the depth zone
where the tr-itium-potential temperature relationship is now linear-) is ve-ry likely valid.
In any case, as can be seen in reconstruction given in figur-e 8-22, the inventor-ies of excess radiocarbon do not critically depend on the exact shape adopted.
In the temperate zones the correlation between t r-i tium and
14 C is not as good as it is in the equatorial zone.
This being
the case, we take another route.
As shown in figure 8-23, theprofile of total 14 C/ C versus depth is remarkably linear- in theseareas of the ocean. This linearity extends to a depth where tritium is about 10% its surface value.
Thus we fall back onto intuition which says that the profile of natural radiocarbon was probably also linear with depth (at least in the zone now contaminatedwith bomb tri tium) .
Again the inventory obtained is rrot highly
sensitive to the choice of this shape (as long as the ~ 4 C values
are required to lie between those for surface water and those forthe base of the tritium-contaminated zone).
Let us now turn to the r-esul ts of this ente r-pr-ise.
The distr-ibution of bomb- 14 C water- column inventories is depicted as afunction of lati tude for the three major oceans in figure 8-24.
The str-iking feature is the several-fold lower inventories forstations in the equator-ial zone than for those in the temper-ate
zone.
A similar- pattern is seen in each of the thr-ee major oceans(only the abbr-eviated north temper-ate zone of the Indian Ocean is

417

+200
PACI Fie AND
ATLANTIC CORAL,-....

0+100

0~.....-•a~

BELIZEFLOR IDAGALAPAGOS

u

v

-<l
0

-100 1925 1930 1935 19401945 1950 1955 1960 1965 1970 1975 1980

DATE OF GROWTH (yrs)
Figure 8-21. fj,14 C values obtained from measurements on corals.
The Belize and Florida corals represent surface water of the northtemperate Atlantic (15,16).
The Galapagos coral comes from anarea of intense upwelling in the eastern equatorial Pacific (17).
The time limit for acceptance of water samples as prenuclear isshown.
A,lso shown are the times of the Atlantic and Pacific
GEOSECS surveys.
These measurements were made by Druffel in thelaboratory of Hans Suess at the University of California San
Diego.

418

14

Figure 8-22. fj. C versus potential temperature and versus depth
for samples taken during the GEDSECS program in the western equatorial Atlantic and in the western equatorial Indian Oceans (opensymbols) (21,23).
Also shown are the tritium values (solid symbols) (51,56).
The linearity between both fj.14 C and tritium concentration with potential temperature in the main oceanic thermocline should be noted.
The prenuclear fj.14 C profile is reconstructed assuming that the natural radiocarbon also showed a linear relationship to potential temperature.
The 14C analyses weremade in the laboratories of Stuiver and Ostlund (21,23).
The tritium analyses were carried out in the laboratory of Ostlund (51,

56) .

419

~14C (%0)
-100

0

~?

~14C (%o)
200

100

0

i

1955-1957

-

K

0
I

100

i

1955-1957

E

E

oX

-100

oX

.4

-.4
J;

..-a...I

~

a...

W

NORTH TE MP.
ATLANT IC
27°N 54°W

0W
0
3 H .... 5 0/0.8

SOUTH TEMP.
ATLANTIC
27°5 37°W~SURF. CONe.

0 . - -......----

E

oX

.4

I

..-a...

wo.8

Figure 8-23- ~14C versus water depth in the temperate zones of theworld ocean for prenuclear samples (open circles) and for GEOSECS
samples (solid circles).
The near lineari ty of the trend forwate~s containing morethan 5% of the tritium concentration insurface wa ter should be noted.
In the case of the Indian Ocean
where prenuclear data is sparse samples collected in 1960 are also
shown (triangles).
The solid lines connecting prenuclear surface
water with the GEOSECS profile for samples having less than 5%
the surface water tritium content represent the reconstructed
prenuclear ~ 14 C profiles.
The GEOSECS 14 C measurements were madein the laboratories of Ostlund and Stuiver (21,22,23).
The
prenuclear and 1960 results are from references 4,5,6,8 and 13.

420

S
LATITUDE
N
Z.t..
O_....._---.,;...........-_...
O_---.,....,.-.;;;;;p.-_...._4..0

I

ATLANTIC
OCEAN

I

I
0\
z

a

co
a:::
<t
u

o"-_

_

24r-----'-_....._ .......-._....._..a-_-'_ _- _ - _ ~ _

o
ON 20
<t E

W INDIAN
OCEAN

a::
~ 16
(J)
(f)

E
w2

(J)

x a')

12

>-

8

U

0

wo

a::

o

I-

o

Z
W

>
Z

o"-_.a-_2 4 r - -....-

0.....I.-'-_-""_---..-_--.j--r--.,....,.-....

PACIFIC

OCEAN

*l:P

*

0 ......_1-_"'-.....1-......._
40
S......._ . 1 -..........._..1-_-'0
LATI TUDE

40
N

Figure 8-24. Water column inventories of bomb- 14 C in the temperate
and eq ua to rial regions of the three rna j or oceans.
The 14 C measurements were carried out in the laboratories of Ostlund andStuiver (21,22,23).
The integrations were performed by the authors (75).

421

anomalous in this regard).

EXPLANATIONS FOR THE LOW-EQUATORIAL BOMB C-14 INVENTORIESThe reason that this large inventory difference comes as asurprise is that the driving force for the input of bomb radiocarbon (i.e., the difference between the 14 e/ e ratio in atmosphericcarbon and surface ocean carbon) is very nearly the same in theequatorial ocean as in the temperate ocean. This can be seen fromfigure e-3.
Thus, if the impedance to gas exchange (i.e., thethickness of the stagnant film) is on the average about the samein the equatorial zone as in the temperate zone, then nearly thesame amount of bomb radiocarbon must have entered the equatorial
ocean as the temperate ocean.
Three explanations can be put forth to explain the difference:
1) It is the result of the combination of rapid mixing along
the isopycnal surfaces and the close spacing of the isopycnal horizons in the equatorial zone.
As the volume of water above any
given isopycnal surface is much smaller in the equatorial zones
than in the temperate zones (see figure 8-5), rapid mixing along
isopycnal horizons would lead to the accumulation of more tracer
in the temperate than in the equatorial zone.
2) The impedance to gas exchange is several times greater inthe equatorial than in the temperate zone.
3) The upwellin~ of water in the equatorial zone purges thisregion of the bomb _1 e invading from the atmosphere.
The upwelling water pushes aside surface waters in the equatorial zone,
driving them laterally into the adjacent temperate zones.
This
process would decrease the bomb _14 e inventory in the equatorial
zones and enhance that in the temperate zones.
Of these explanations, we will show that the first can beruled out;_ the second is unlikely but remains a dark horse; andthe third, while quite plausible, is not without problems.
The first explanation can be ruled out because the gradientin bomb radiocarbon along every isopycnal and every horizontal
surface down to the base of the tritium-contaminated zone istoward the equatorial zone.
Horizontal or isopycnal mixing would
increase rather than decrease the inventory of bomb - 14 e in theequatorial zone.
Hence the bomb radiocarbon entering the equatorial zone from the atmosphere cannot be dissipated by such mixing.
This is a firm argument; the first explanation must be crossed
from the list!
In chapter 3 we discussed the factors controlling the rate ofgas
exchange between the ocean and atmosphere and concluded thatthe dominant factor must be wind speed.
From meteorological observations it can be shown that the average wind veloci ty in theequatorial zones is about 85% that in the temperate zones.
Also,
the average stagnant film thickness as measured by the radon method in the equatorial zone is only about 20% thicker than thatmeasured in the temperate zone (176). Hence both theory and measurements would suggest that the gas exchange impedance in theequatorial zone is perhaps 1.2 that in the temperate zones.
To
explain the difference in inventories the impedance in the equa-422torial zone would have to be three to four times that in thetemperate zone!
While this appears quite unlikely, this explanation remains a dark horse.
Evidence for upwelLing along the equator and along continental margins in the equatorial zone comes from every discipline.
The geologist sees the products of upwelling (for example, thefrustule of diatoms) in enhanced abundance in the sediments accumulating beneath these zones, the biologist observes enhanced
rates of plant growth in these zones, the chemist observes highconcentrations of P0 4 , N0 3 , H 4 Si0 4 , C02, ... in the surface watersof these zones, and the phys ical oceanographer predicts such upwelling based on the patterns of winds.
Thus few would argue
against upwelling.
The questions in connection with the radiocarbon inventories have to do with the total flux of upwelled water(is it great enough to purge the equatorial zone of its radiocarbon on the proper time scale?) and with the origin of the upwelling water (is it recycled equatorial water or does it come fromthe adjacent temperate zones or from intermediate waters flowing
beneath the equatorial zone?). Other disciplines have little firm
evidence to offer in answer to these questions.
They are difficult ones.
This being the case, we must try to get this information fromthe distributions of the tracers we have already mentioned (C02,
3H and14 C) and perhaps
supplement this with information from~racers whose potential in this regard has yet to be tested (3 He ,5 Kr , ..• ).
Let us start by considering bomb radiocarbon alone.
To make this calculation as straightforward as possible, we will
make six simplifying assumptions.
One looks for the big things
first; the complicating details can be added later.
First, wewill assume that the upwelling water carries no bomb _14 C (thisassumption leads to a lower limi t on the flux).
SeconC\, we willassume that at the time of the GEOSECS surveys the bomb- 4C· budget
in the equatorial oceans had reached steady state (i.e., the inputof new radiocarbon from the atmosphere nearly balanced the loss
resulting from the purging action of upwelling water).
This
assumption is supported by the 14 C measurements on corals (seefigure 8-21).
The third assumption is that there is no input ofbomb-produced 14 C via lateral miXing with the adjacent temperatezones.
Like the first assumption, this one leads to a lower limiton the upwelling rate. The fourth assumption is that mixing along
isopycnal horizons within the equatorial zone is so rapid that theexact route taken by the upwelling water has no consequence.
As
shown in figure 8-22, the tracers are well mixed along isothermal(and hence also isopycnal) surfaces-within the equatorial zone.
This assumption allows the calculation to be carried out as if
upwelling were uniform throughout the zone (which it clearly isnot).
The fifth assumption is that the C02 partial pressure insurface water is the same as that in the air.
The sixth assumption is that the rain of bomb radiocarbon with falling particles
is negligible.
While nei ther of the latter two assumptions isstrictly correct, only second-order changes in the calculated
fluxes occur when they are handled properly.
The calculation then becomes one of balancing the water andradiocarbon budgets for the upper equatorial zone. The fluxes are423

Figure 8-25. Simplified model for the bomb _14 C balance in thesurface mixed layer of the equatorial ocean.
Water upwelling into
this layer is dissipated by poleward flow along the surface.
~hile in residence at the surface,
the upwelled water gains bomb
4 C through C02 exchange with the atmosphere.

EVASION

11

INVASION~

51 DEWELLJ NG",..
SOUTH..,,-.-_-.....--...,. '·SIDEWELLING

U

NORTH
SURF.
MIXED
LAYER

RAIN OFPARTICULATEDEBRISUPWELLING(ASSUMED TOCARRY NOBOMS I4 C)

Table 8-5- Summary of the data parameters used to calculate thearea averaged upwelling rate in the equatorial ocean from the distribution of bomb-litC.
(The asterisk behind 14 C designates
excess over natural).
GEOSECS~litC

PRENUCLEAR
~14C

~14C

14 C*/C)atm

atm

sw

14 C */C)sw

0/ 00

0/ 00

0/ 00

440

95

-25

-65

2.9

15

1978

330

90

-25

-65

2.3

11

1974

390

80

-25

-65

2.9

15

t\14 C

Date of
GEOSECS
Survey

atm

sw

0/00

Atlantic

1972

IndianPacific

424

W
m/yr

shown in figure 8-25.
The re are just two wa ter fluxes.
Loss tothe temperate zones along the surface must balance upwelling.
There are three fluxes involved in bomb- 14 C balance; the invasionof bomb- 14C bearing atmospheric C02, the evasion of bomb-l~ bearing surface ocean CO 2 and the loss of bomb- 14 C with the water exported from the equatorial zone to the tem~erate zone. As the upwelled water is assumed to carry no bomb- l C this source does notcontribute to the 14 C budget.
The

14C

balance can then be written as follows:l~C*)atm 1 = l~C* )sw [1 + WCl:C02JsW]

8-1

where I is the rate of C02 exchange (~ 16 moles/m 2yr), [L:C02Jsw
the concentration of inorganic carbon in the exported surface
water (~2 moles/m 3 ), 14 C*/C )atm and 14 C*/C )sw are respectively
the bomb _14 C/ C ratio in the atmospheric CO 2 and in surface water
L:C02, and W the unknown upwelling rate (in units of m 3 /m 2yr orm/yr) averaged over the entire area of the equatorial zone. Thisequation can be solved for W yielding:8-2

As can be seen from table 8-5, all three oceans yield an area
averaged upwelling rate on the order of 13 m/yr.
Since the aggregate area of the equatorial ocean (15°N to 15°S) is about 1 x
10 14 m2 the total upwelling flux is about 1.3 x 10 15 m3/ yr (- 40
Sverdrups).
Thus the amount is about the same as the flux of new
bottom water to the deep sea.* Due to the assumption that the upwelling wa-ter contained no bomb radiocarbon and the assumptionthat no lateral or isopycnal mixing takes place across the 15°N
and 15°S boundary zones, this flux must be considered a minimum.

IMPLICATIONS OF EQUATORIAL UPWELLING TO THE TRITIUM BUDGETThis result has an important implica tion to the budget oftritium in the ocean.
If the water upwelling into the equatorialzone is free of bomb- 14 C, it must also be free of tritium.
As at
the time of the GEOSECS surveys the input tritium from atmosphere(and rivers) to the ocean had dropped to a very low level, upwelling should have been purging the tritium present in the equatorial zone (by pushing it along the surface into the adjacent temperate zones).
The approximate time required for half the tritium
to be purged is given by the following equation:*We do not wish to imply that these two fluxes need be related in
the sense that equatorial upwell ing is the route by which deep
water returns to the surface, but rather we use the comparison asa means of putting this very large flux into context.

425

tl/2 = 0.693 h/W
where h is the mean penetration depth of tritium in the equatorialzone (-180 meters) and W is the rate of upwelling (at least 10
meters per year).
So we get a half-time for the purging of tritium no more than 12 years.
This corresponds to a purging loss ofat least 8% each year.
Is there evidence for such a decrease?
There are a few places in the equatorial ocean where thetri tium profiles measured as part of the GEOSECS program can becompared wi th those measured at earlier times.
The changes ininventory found in this way are summarized in table 8-3. Althoughscanty, and sUbject to rather large uncertainties these results
suggest that the inventories held their own and perhaps evenincreased somewhat during a several-year period preceding the
GEOSECS surveys.
The validity of this apparent contradiction to the upwellingmodel can be checked by a less direct method.
No means exists bywhich a decrease in the mean penetra tion depth of tri tium wi thtime could occur once the major amount of tritium had been delivered to the sea.
Hence, if the water column inventory of tritiumin the equatorial zone decreased, then the surface water concentration must also have decreased.
If the mean penetration depth
were increasing the surface water concentration would decrease at
an even higher rate than the water column inventory.
In the western equatorial Pacific Ocean, where we have both extensive tritium
measurements on surface water and a sequence of 90 Sr measurements
on a coral head, the surface water shows a 3% per year in~bease intritium content (see table 8-4) and almost no change in
Sr content (see figure 8-18) between 1968 and 1973.
Thus we have aproblem.
The upwelling hypothesis, which so neatly explains thebomb- 14C distribution, appears to run into a snag for tritium.
One way around this difficulty (at least for the Atlantic andPacific) is to bring tritium into the equatorial zone from thelarge reservoirs of high-tritium water in the north temperatezone. This transport could be accomplished either by lateral mixing across the 15°N front or by advection of upper temperate thermocline waters into equatorial zone (as a contribution to thetotal upwelling flux). As the north temperate water which carries
large amounts of tritium also carries bomb radiocarbon, the addition of this radiocarbon to the equatorial zone would then have to
be balanced by an even greater rate of upwelling.
Further, as wehave shown, the 8-S relationships in the equatorial zones of theAtlantic arid Pacific are more akin to those of the southern temperate zone than to those of the northern temperate zone.
Thus ifnorth temperate thermocline waters are brought in to maintain thetritium inventory then waters from the south temperate thermocline
would also have to be brought in to maintain the proper 8-S relationship.
The south temperate thermocline waters would also carrybomb- 14C (but little tritium).
Again more upwelling would be required.
While it is possible to create a blend of all thesewaters which reproduce the distributions of bomb- 14 C, of tritium,of temperature and of salinity, the total rate of upwelling reqUired goes up by at least a factor of three.

426

Could the total flux of upwelled water in the equatorial
zones of the world ocean be as high as 120 Sverdrups?
While nofirm grounds for rejection can be given, many physical oceanographers would be very uneasy wi th such high fluxes.
Thus someother new information is needed to tip the balance toward or awayfrom acceptance of the upwelling explanation for the low equatorial zone water column inventories of bomb radiocarbon.

AN UPWELLING RATE BASED ON THE EQUATORIAL C02 ANOMALY
We laid the groundwork for another upwelling rate estimate bydiscussing the distribution of chemical properties in the surface
waters of the equatorial zone in the Pacific Ocean (see chapter3). It was shown that the concentration of virtually every property is anomalous compared to the adjacent surface waters.
The
temperature and salinity are low and the nutrient element concentrations and C02 partial pressure are high.
It is the high value
for the PC02 which permits an estimate of the upwelling rate.
Were new CO 2-laden water not entering the surface layer, the CO 2
anomaly would be dissipated by loss to the atmosphere.
In orderto obtain a quantitative estimate of the rate of this renewal, theamount of CO 2 lost to the atmosphere by equatorial surface waters
and the rate of CO 2 evasion to the atmosphere must also be known.
It is the first of these that is not so easily obtained. Seemingly it would require that not only the ~C02 content of upwelling
water be known but also the amount of C02 lost to plants.
As theamount of ~C02 in the upwelling water depends on the depths fromwhich these waters originate (an unknown), we are seemingly stuck.
As already discussed in chapter 3, there is a clever way to circumvent the need to define the upwelling depth.
This is accomplished using plots of ~C02 (normalized to a constant salinity)
against NO 3 for samples from the thermocline and surface.
Plotsof this type for three la ti tude zones are shown in figure 8-26.
In each case the points for thermocline waters fall along a trendwith a slope close to the Redfield ratio.
Thus the photosynthesis-respiration cycle explains the variations seen.
For samples taken from the uppermost thermocline and at the surface thepoints fall below this trend.
The likely explanation is that the~CD2 content of these waters has been al tered by CO 2 loss to theatmosphere.
If so, then we can take the magni tude of the deviation of the points for the surface water samples from the trend
es tablished for thermocline samples as the amount of CD 2 lost to
the atmosphere.
Surface water samples from the 4°N to 4°S equatorial strip
showing high PCOZ values have an average ~C02
deficiency of 60 lJm/kg.
Each unlt of upwelled water appears tohave lost this amount of C02 to the atmosphere.
The balance canbe written as follows:~~~

I [ PC 02 S r f _ 1J = W C02
PC0 2alr

8-3

where PC02surf and PCO air are mean CO 2 partial pressures
for surface waters and ai~ in the equatorial zone, I is the inva-

427

2300

2300

CI

ONSET

..ll:

.......
0

DEPARTURE
FROM RBJAELD

2200

2200

SLO~

E

*

o

10

20

40

30

~

_

+

-

-

_

1900 "• -_ _-+--_ _-L-_ _

-

~

-

2000

l

~ 2000-2100-::t. 2100

I-+._ _"""-_ _-1-..-..

. 1900 J.._ _-.l.
100(N03)* (p.mol/kg)

20

30

40

(N0 3 )* (J,Lmol/kg)

2300
CI

..ll:

.......

2200

0

E

3-

*N
0

2100
2000

U

t!
1900
10

0

20

30

40

(N0 3 )* (p.mollkg)

.

... .
~~

2300
CI

.x

.......

<5

2200

..

.-' ~

2300

2200

E

::l

2100

2100

2000

2000

*N
0

U

W

1900
0

19 00 l-.-_ _~_ _--I-_ ___J__ __ _ L _ _....
10

20

30

40

o

10

20

(N0 3 )

(N03)* (fLmol/kg)

30* (fLmol/~g)40*' NORMALIZED TO A CONSTANT SALl NITY
Figure 8-26. Plots of LC02 (normalized to a constant salinity)
agains t ni t ra te for sample s of the surface and the rmocline wa te r
in five latitude belts in the equatorial Pacific (at longitude
155°W).
In each case a straight line is fit to the points withintermediate nitrate content.
As these line have slopes close to
the Redfield ratio (~6.8) it is assumed that these samples havemainly been influenced by the photosynthesis-respiration cycle.
For shallower waters (i.e., those with lower N03 concentrations)a deviation from this trend is seen. This deviation is presumably
due to C02 loss to the atmosphere.
For deep waters (i.e., thosewith higher N0 3 concentrations) a deviation is also seen.
These
deviations
mustreflect
calcitedissolution
as
its
onset
corresponds to the depth where the alkalini ty begins to rise.
These results were obtained as part of the Hawaii-Tahiti Shuttle
program (433). This is a repeat of figure 3-26.

428

sion ra te of CO 2, W is the unknown upwelling rate, and ~ L:CO 2 isthe amount of CO 2 lost by the equa torial surface waters.
Theequation states that in order to maintain these distributions as
they are, the C02 being lost to the atmosphere must be replaced by
the addition of upwelling water.
Solving for W, we get:w

I

PCO 2 s~rf _ 1]
[ PC0 alr2

8-4

In the equatorial Pacific (4° N to 4° s) the PCO 2 in the surface
water averages 410 x 10- 6 atm, while that in the air is 330 x 10- 6
atm.
Thus the term in parentheses is aqout 0.25.
I is about 16
moles/m 2 yr.
bL:C02 is about 0.06 moles/m.
Thus:W

16 moles/m 2yr
0.06 moles/m 3

x 0.25= 66 m/yr

To be compared with the bomb radiocarbon-based rate which applies
to a latitude belt 15°N to 15°S, this estimate must be reduced toabout a factor of 4/15 (hence from 66 to 18 m/yr).
Thus both the
distribution of bomb _14 C and the chemical anomalies of carbon in
the equatorial zone lead us roughly to the same conclusion.
Thearea averaged upwelling rate in the equatorial zone lies in therange 16 to 40 m/yr and amounts to a global flux somewhere in therange 50 to 120 Sverdrups.
We are forced to express our results as area-averaged upwelling rates because the rapidly flowing equatorial currents areable to carry the upwelled water from one "end" of the zone to theother during their time of residence at the equatorial surface.
Because 14C requires a longer time to equilibrate with the atmosphere (-15 years) than the time of residence of water at the surface of the equatorial zone (-6 years) its distribution has lost
any message carried by the geographic distribution of upwelling.
CO 2, on the other hand, has an equilibration time (-1 year) shorter than the residence of water at the surface (-6 years).
Thus
the distribution of the PC02 anomaly carries information regarding the point of entry of the upwelling water.
It demands thatthe upwelling occur both along the equatorial belt and in coastal
upwelling areas from which currents feed into the equatorial zone.
As shown in the maps for the equatorial Pacific, Atlantic andIndian Ocean (see figure 8-27), only in the Pacific has a prominent surface water CO 2 partial pressure anomaly been found along
the equator.
Where then does the upwelling occur in the Atlantic
and Indian Oceans?
Al though we do not know the answer to thisquestion our suspicion is that in all three oceans coastal upwelling makes a very important contribution to the total.
As can beseen from the map for the equatorial Pacific Ocean, the region
south of the equator along the South American coast appears tocontribute strongly to the anomaly.
The signature of thisupwelled water is apparently carried far from the continental margin by the prevailing currents.
A similar area of strong upwelling occurs along the coast of Africa to the south of the equator.

429

ATLANTIC

INDIANPAC I F Ie
Figure 8-27. Maps showing the dis tribu tion of excess CO 2 partial
pre s sur e s in eq uat 0 ria 1 sur fa c e wa t e r .
Th e con t ou r s are in un itsof 10- 6 atmospheres.
These maps were constructed by Taro
Takahashi of the Lamont-Doherty Geological Observatory using datafrom many expeditions
run in different years and seasons(387,388,393,394,406).

430

ATLANTIC

INDIANPACIFIC
Figure 8-28. Radiocarbon measurements (expressed as ~14C) on surface water samples from the equatorial zone.
The arrows represent
the paths followed by waters upwelled into coastal areas.
The
results were obtained as part of the GEOSECS program (21,22,23)
and as part of the Hawaii-Tahiti Shuttle Program (Quay andStuiver personal communication).
The results obtained by Ellen
Druffel on a coral from the Galapagos are also included (17).

431

We suspect that this area is the conduit for most of the upwelling
water reaching the Atlantic equatorial zone.
In the Indian Ocean
strong upwelling occurs in the Gulf of Arabia.
Again this may
be the major conduit for upwelled water reaching the equatorialzone.
The radiocarbon content of surface waters in and near thesezones of intense upwelling shows a local anomaly (exceeding thatseen in the ambient surface waters of the equatorial zones).
The
map in figure 8-28 shows the results for surface waters collected
between 1972 and 1980 in the equatorial zones of the world o~~an.
As shown by the resul ts on corals (see figure 8-21), theC/Cratio in surface water was quite uniform over this time interval;
hence despi te the differing times of collection the resul ts onthese samples should be directly comparable.
It can be seen thatthe lowes t values are found in and downs tream from the areas ofintense coastal upwelling (namely, off South America in the eastern Pacific, off Africa in the eastern Atlantic, and in the Gulfof Arabia in the northwestern Indian Ocean).
Unfortunately, we do
not as yet have systematic C02 partial pressure surveys for thesurface waters in these areas.
Also, the 14 C data are quite
sparse.
HE-3 DISTRffiUTION IN THE MAIN OCEANIC THERMOCLINEIn the preceding sections we have provided evidence supporting a strong role for upwelling in the ventilation of the equatorial ocean. This upwelled water must flow along the surface into
the adjacent temperate zones.
Then where does it go?
Some of it
flows across the temperate zone into the polar areas and contributes to new deep water formation.
Some sinks within thetemperate zone.
This sinking is likely via flow along isopycnal
horizons from thei r winter ou tcrops.
Tha t portion of the water
entering the temperate thermocline eventually passes across the
15° fronts into the equatorial zone completing the loop.
Yet another of our tracers gi ves us some insight into theprocesses responsible for the ventilation of the temperate thermocline.
The 3 He generated by the decay of bomb tritium adds
important information not obtainable from the tritium distribution
alone.
In order to appreciate the added information provided by 3 He
let us first consider the vertical distributions of t ri tium andbomb- 14 C w~th depth in the equatorial and temperate zones.
Depth
profiles for these isotopes in the equatorial and temperate zoneof the North Atlantic are shown in figures 8-29 and 8-30.
The
simplest model to explain these distributions would be one involving upwelling in the equatorial ocean coupled with downwelling in
the temperate zones.
Half of the water upwelling in the equatorial zone is assumed to flow across the surface to the north andthen to downwell in the north temperate zone.
The other naIf goesto the south temperate zone.
In both zones eddy diffusive mixing
in the vertical is also assumed to operate. As the density gradient in the equatorial zone is much larger than that in the temperate zone, the eddy diffusivities need not be the same in the twozones.
A mixed layer 100 meters thick is assumed in the equator-

432

14

E XCE sse (0/00)

o0;...._ _~4,;.0_ _--.;;;.8.,;.O_ _......;12r-0

16r-0-r:'_ _2..,00
I

8

1

I
/

200
~/

y/

.........

E 400
I

//

I-

~

//

/

// 8

eL
W 600

/

/

IA

J

/

/

EQUATORIAL (0)2

K=o.35 em Is
W= 16.0 ml yr
E = 19.0 mol/m 2/yr

/

Cl

//8

TEMPERATE (8.)/

K =2.8 em 2 /s
W= -5.7 m/yr

4/
800

I

E =22.0 mol/m 2/yr

8'
I

I
I
1000 .....- -........- - -........- - - ' - - - - - - - - - -...

Figure 8-29. Mean profiles for !J.14 C in the north temperate andequatorial zones of the Atlantic as of 1973 (21) compared with theprofiles (lines) yielded by the regional 1-D model described in
the text.
The parameters used in the model are given.
Also shown
are the thicknesses assumed for the surface layers in theseregions.

433

3 H (tu)3 H (tu)
0.0
0

0.8

0.0
0

3.2

E
240

600

w

EQUATORIAL

400

480

6.0

8.0

•
N. TEMPERATE

2

2

K = 0.35 cm Is
W= 16.0 m/yr

320

4.0

400

Cl.

a

2.0

200

-

160

~

2.4

,-,A

-.

80

I

1.6

800

K=2.80cm /s
W= -5.7 m/yr

1000

-

1200

Figure 8-30. Mean profiles of tri tium in the north temperate
eq ua torial zones of the Atlantic Ocean (52,52) compared wi th
profiles (lines) yielded by the regionalized I-D model.
parameters used in the modeling are those for the best fit tobomb_ 14 C distribution (see figure 8-29).

434

and
the
The
the

ial zone and a mixed layer 150 meters thick is assumed for thetemperate zone.
The latter is larger because of deep winter convection which thickens the mixed layer in the temperate zone.
The
CO 2 invasion rate is taken to be 19 moles/m 2 yr in the equatorialzone and 22 moles/m 2 yr in the temperate zone (consistent with thedifference in rates based on wind velocity and on radon measurements) .
A numerical model is then constructed which allows ~hese processes to be simulated.
The model runs begin with the
4 C at its
natural steady state and with no tritium. Bomb tritium and radiocarbon are then added to the sea surface in accord with the temporal and geographic input functions described above. The upwelling rate in the equatorial zone and the vertical eddy diffusivities in both zones are then adjusted to give the best fit.
The
results are shown in figures 8-29 and 8-30.
As can be seen, thedistributions can be satisfactorily explained by this simple linked regional l-D model.
The equatorial upwelling rate obtained inthis way is similar to that calculated above.
The eddy diffusivity needed for the temperate zone is about 8 times higher thanthat needed in the equatorial zone.
Both the radiocarbon andtritium distributions are fit reasonably well.*
The fact that the model fits cannot, of course, be taken as
evidence that the ocean functions in accord wi th the model.
Theability of the simplest of models to fit the primary features ofthese distributions rather says that these distributions carryli ttle informa tion about the na ture of the ventilation process.
Models involving ventilation along isopycnals could equally wellachieve the observed distribution.
Wi th this background in mind, let us consider the distribution of 3He .
The properties of this tracer of importance to the
problem of thermocline ventilation are, first, that its production
is tied to the radiodecay of tritium; and second, that when thermocline water carrying excess 3He reaches the surface, its excess
3He is lost to the atmosphere.
Thus, unlike the other anthropogenic tracers which are produced in the atmosphere and from there
penetrate into the sea, 3He is generated within the sea and escapes to the air.
In the absence of mixing wi thin the sea
the
3He/ 3H ra tio would cons ti tu te an hourglass type keeper of time.
Once isolated from the air, a Nater samp\e w~uld accumulate the3He produced by the decay of 3H and theHe/ H ratio would be asimple function of time; the relationship being:

8-5
where 3He* is the concentration of excess mass three helium inthe sample and 3 H is the concentration of tritium in the sample.
This ideal situation never exists in the ocean.
All the samples
we obtain are mixtures of waters which left the surface at*The fit for tritium in the equatorial zone would be better if an
input rate 25% higher than that obtained by the procedures outlined in this chapter were used.

435

24.0

24.0

•.,

•

••

~ ATMOSPHERICSATURATION'0 25.0

~

>-

~

(f)

Z
W

a 26.0-J«t-z,~ •••
I

W

l0

a.

,
I
I

I
27.0

I

•

>-

t-

(f)

Z

••• •

W

o 26.0-

• •••
••

~lJ
..., ..

i=
z
w

••
•
•
•

l-

~ 27.0-

100

.

~

-J

~

I
I
28.0

-

~ 25.0 I-

28.0

.,.

.~

••

•••

•

,.- "

-

I

5

0

3H (tu)Figure 8-31. Plots of 3 H (in T.U.) and excess 3 He (in C;He units)
as a function of density for samples from the upper 1000 meters of
the water column in the Sargasso Sea.
Densi ty rather than depth
is used to permit the overlay of results taken at different times
and places.
The important point these results make is that there
is a sharp break in the profiles at a density of about 26.4 0 /00.
Above the density horizon the tritium is well mixed and the waters
are free of excess 3 He ; (the difference in solubili ty of 3He and
4 He leads to a C;3 He value of -140/00 for surface water at equilibrium wi,th the air).
At a density of about 26.4%0 the excess 3 He concentration shows a sharp increase.
For more densewaters both the concentrations of tritium and of excess 3 He decrease with depth to a density of about 27.2%0.
These measurements were made by Bill Jenkins of the Woods Hole Oceanographic
Institution (279). See page 681 for definition of C;3 He •

436

100

80

•
,.....,.

•

-

60

••
2.0
•

•

.........
::l1.5+.........,

0

<V

0~
.........,
(1)

J:

40

I
rt)

to

20

rt")

•

••

•

W
0

w

....-,
•
2

(f)
(f)

X

0

-20
0

1.0

3
3 H (tU)

4

5

0.5

0.0
6

Figure 8-32. Plot of o3 He versus tri tium concentration for thedata in figure 8-31.
The curve obtained from the l-D model of
figures 8-~9 and 8-30 is shown for comparison.
No l-D model canyield the distributions of both these tracers.
Lateral processes
must be operative.

437

different times carry~ng different tri tium concentrations.
Thus
we can only interpret He data in the context of some sort of mixing model.
Shown in figure 8-31 are profiles of 3H versus potential
density and of excess 3He versus potential density for profiles
taken in the Sargasso Sea of the North Atlantic by Jenkins ofWoods Hole.
Potential density rather than depth is used so thatdata from different stations and different times can be shown
together. *As can be seen, the upper horizons (constituting
approximately the 0 to 400-meter depth range) has nearly constanttri tium and is nearly free of excess 3He .
This water must be
returned to the surface every winter and lose the 3He generated
during the previous year.
Below this layer the tritium content
begins to decline and the excess 3He rises abruptly to a maximum
and then declines along with the tritium. A secondary-maximum inboth tri tium and 3He is seen at a potential densi t~ of abou t
27.5 0 /00.
Beneath this, both the tritium and excess He content
drop to zero.
Jenkins showed that, even if the vertical diffusivity wereallowed to change with depth, no 1-D model could explain the dist ributions of these isotopes for densi ties less than 27.5 0 /00.
(Of course it goes without saying that the secondary maximum could
not be generated by vertical processes alone.)Thus the 3He data
provides a compelling confirmation that ventilation along isopycnal horizons must be very important in the temperate regions!
Jus t to show how bad the fi t would be using the 1-D model
discussed above, we have compared in figure 8-32 the excess 3 He
versus 3H curve based on the measurements with that calculated by
the I-D model.
This model (or any I-D model) cannot produce thelarge bulge to high 3He and tritium values shown by the data!
PURPOSEFUL TRACERS

So far we have discussed only tracers which happen to be present in the sea.
While extremely valuable to the field of oceanography, these tracers appear to be incapable of yielding morethan an upper limit on the rate of diapycnal mixing (i.e. thoseyielded by the l-D model).
As the entry of the anthropogenic
tracers into the main ocean thermocline can be explained entirelyby along-isopycnal penetration from surface outcrops, it is possible that little diapycnal mixing occurs within the main oceanic
thermocline.
Also, the distributions of anthropogenic tracer do
not require any across-isopycnal advection in the temperate zones.
Is it possible that tracers could be introduced into the sea by
the oceanographer in such a way that the time evolution of theirdistributions would prove more enlightening in this regard?
A tracer (some perfluorocarbon not currently present in nature)would be injected at a single point in the main thermocline.
Careful attention would be given to keeping the initial patch asthin as possible and to documenting the densi ty horizon on which
it came to res t.
Once in the ocean, the tracer would be allowed*Internal waves within the thermocline can cause temporary changesof the height of a given isopycnal by several tens of meters.

438

to spread for several years before an attempt was made to locate
or map it.
A search would then be launched.
If it could befound, a program of mapping would be ini tiated.
Ideally, then,its vertical distribution would be determined as a measure of thedegree of cross isopycnal mixing, the displacement of the mediandensity horizon would be determined as a measure of the extent of
upwelling or downwelling, and its lateral dimensions would be determined as a measure of along-isopycnal mixing.
One might ask whether the spread of such a tracer in the vertical could be the result of contact with the outcrop during thewanderings of the patch along its isopycnal horizon.
The answer
would be no.
The reason is that if the tracer did reach the outcrop, it would be lost to the atmosphere!
Hence, any cross isopycnal mixing observed during the experiment would reflect only
mixing processes within the interior (including frictional mixing
along the "sides" of the basin).
An even more elegant experiment could be carried out if two
tracers were added together: one a gas and one a nonvolatile compound.
The nonvolatile compound would not be lost during outcrop
contacts and hence, if along-surface mixing followed by reentry
into the thermocline were to occur, it would be recorded by the
nonvolatile compound.
Is it feasible to do such an experiment?
The answer is yes.
The reason is that gas chromatographs are capable of detecting as
little as 10 6 molecules of various perfluorocarbons. If, for example, 30 kilograms of a perfluorocarbon compound (molecular weight"'300 g/mole) were added to the sea, there would be 6xlO 2 5 molecules to trace. Thus even after the tracer had been spread over %volume 50 00 kIn on a side and 100 meters thick (i. e., 2. 5xl 0 1
liters) there would still be 2.4xl0 7 molecules per liter of water.
Perfluorocarbons are biologically inert.
Hence the tracer
would not be degraded by marine organisms nor would its addi tionpose any pollution hazard.
The large amount needed for the experiment could be diffused into the ocean through silicon tubing thus
avoiding the need to predissolve it in many tons of sea water.
A
nonvolatile compound could be made by SUbstituting an ionic groupin the structure of a perfluorocarbon.
Analysis of this compound
could then be done by fluorinating the sample and converting the
nonvolatile back to a volatile compound.
One very serious problem could void the whole concept. It isalmost certain that the initial patch would break up into a seriesof filaments.
No one knows on what time scale these filaments
would coalesce into a continuous patch.
Were the patch to break
up into a multitude of fragments and remain even years later after
dispersal to all parts of the isopycnal horizon in this inhomogeneous state, then mapping to yield the type of information weseek would prove impossible.
Once again the ocean would havefoiled an attempt by man to pin down the details of its operation!
SUMMARY

The distributions within theseaofthe anthropogenic
tracers, bomb tritium and bomb radiocarbon, have much to tell
about the ventilation of the upper ocean and about the formation

439

EQUATOR

Figure 8-33. The authors' guess regarding the pathways followed by
the water ventilating the main oceanic thermocline.

440

of deep water.
The first global survey of these distributions wasmade as part of the GEOSECS program.
While clearly showing wherematerials added to the sea surface go on the time scale of ten
years, these results provide only hints regarding the pathways ofthis penetration.
The hints with regard to the pathways followed
by water ventilating the thermocline obtained from bomb _14 C andbomb - 3 H must be tested through the use of other data.
For example, the large C02 partial pressure anomaly in the surface
waters of the equatorial Pacific can only be maintained by rates
of upwelling somewhat greater than the minimum value required to
explain the distribution of bomb radiocarbon.
Another example isthe use of 3He to show that the distribution of tritium in thetemperate thermocline cannot be the result of downward mixing
across isopycnals.
It must rather involve mainly movements along
isoPlcnal surfaces. Further clues await us from the distributionsofHe and of 85 Kr in the oceanic thermocline.
The purposefuladdition of man-made tracers within the thermocline would greatly
facilitate our understanding of these phenomena.
What may emerge from these studies is the pattern of ventilation shown in figure 8-33.
Water descends into the main thermocline from the outcrops of isopycnals in the temperate zones.
This water leaves the thermocline by a combination of downwellingfrom the temperate thermocline into the underlying intermediatewater and upwelling in the equatorial zone. The flux of temperatesurface water upward through the equatorial thermocline is supplemented by water coming from the intermediate waters beneath thethermocline.
The water reaching the surface by upwelling spreads
to the north and south along the surface, returning to the temperate zones for another cycle.
PROBLEMS
B-1

What is the value of the apparent coefficient of verticaleddy Diffusion in a water column which has an upwelling rateof 10m/yr and z* value of 250 meters?

e-2

If a near-bottom water sample from the North Atlantic collected in 1976 yields an excess 3 He to 3 H ratio of 0.8,
(in atoms)
when did this water leave the surface?
Assumethat this water is free of mantle-derived helium and that it
did not mix with other waters.

8-3

the time of the Atlantic and Pacific GEOSECS surveys, thevalue for atmospheric CO 2 was close to 400 0 /00.
How
many excess 14 C atoms were there in the atmosphere? How many
were in the Ii vin~ biosphere?
The mass of carbon in theatmosphere is 6x10 5 moles.
That in the living biosphere isalso about 6x10 15 moles.
Most of that in the living biosphere is in trees.
State your assumptions.
The specificactivity of carbon with a l:!.14 C of 0% is 13.0 dpm/g.
A~

l:!.

8-4

4C

The production rate of tritium by cosmic rays is about 0.2
atoms/cm 2 sec.
If prior to bomb testing the 85% of the tritium produced reached the oceans before undergoing radiodecay,
and if the mean mixing depth of this tritium was 350 meters,then what must have been the average tritium concentration in441

surface ocean wa ter (in T. U. )?
One T. U. corresponds to one
tritium atom per 10 18 hydrogen atoms.
The half-life of tritium is 12.4 years.
8-5

During the summer the mixed layer depth in the temperate
ocean is about 50 meters.
Winter cooling causes it to
thicken to about 150 meters.
During the early 1970' s, when
the atmosphere had a L\14 C value of about 400 0 /00 and thesurface ocean a value of about !~o%o, what sort of seasonal cycle would one expect in L\ C?
Assume a ~C02 concentra tion of 2.0 moles/m 3 and an exchange rate of 20 moles/
m2yr.
For simplicity, also assume that the mixed layer has
the summer depth for six months and the winter depth for six
months.
Also assume that during the early 1970' s the annual
ave rage L\ 14 C in the surface ocean was cons tant (i. e. , theannual amount of bomb- 14 C gained by COz through the invasionof atmospheric COz is lost to deeper water by below 150
meters vertical mixing).

8-6

If 100 kg of a perfluorocarbon with a molecular weight of 300
g/mole were dissolved in sea water, how many molecules would
be present?
If this tagged sea water were then mixed into a
volume 5000 km on a side and 100 m thick, how many molecules
would be present in a typical one liter sample from this volume?

#8
A geochemist named Lyle de Vendra becomes interested in upwelling and decides to construct a computer aquarium to study theprocesses occurring in a surface mixed layer perched on top of an
upwelling water column.
He envisions a column 1000 meters high.
Water is added to the base of the colunm and upwells at the rateof 15 meters per year.
The upper 75 meters of the column is well
mixed. Beneath the mixed layer mixing occurs so as to produce anapparent eddy diffusivity of 0.5 cmz/sec.
The upwelling water has
a temperature of 10 c C and a salinity of 34.50 0 /00.
Professor de
Vendra first considers the water budget in the mixed layer.
The
rates of rainfall and evaporation are respectively 90 cm/yr and 75
em/yr.
What is the surface salini ty?
The rate of tri tium rain(due to a continuous release from the world's nuclear reactors) is
10 atoms/cm z sec.
What is the steady state tritium concentrationin the mixed layer (in T.U.)? What is the steady state profile oftritium beneath the mixed layer? How would the steady state tritium content of the mixed layer change if the eddy diffusivity beneath the mixed layer were reduced to zero?
Professor de Vendra then turns his attention to the carbon
cycle.
He starts with an abiotic water column.
The water entering the bottom has a ~COz concentration ofa moles/m 3 and apotential * C02 partial pressure of 600 x 10- atmospheres.
The
surface water temperature is 25°C.
The COz invasion rate is 20zmoles/myr.
What is the steady state PC0
for surface water?
2
What is the steady state ~COz?
Assume that dPCOz/pco z = 10
SUPERPROBLEMt.

*By potential PCO

is meant the C02 pressure achieved if the water
18 isochemically aepre8surized to one atmosphere and warmed to 25°C.

442

dECOZ/ECOZ and that this value is independent of PCOz·
14
Getting the hang of things Prof. de Vendra moves on toC.
He assumes that a nuclear society maintains the ~14C of the atmospheric CO~ at +300 0 /00. He assigns the carbon in the
upwellingwater a ~ 4C of -100 0 /00. What would be the steady state ~ 14C
value for the surface mixed layer? Would the answer be any different if he dropped vertical diffusivity beneath the mixed layer tozero? Compare the fraction of PCOz equilibrium with the fEmosphere achieved by the surface water wi th the fraction of
C/Cequilibrium.
Gaining momentum Prof. de Vendra decides to put biology into
his model. He inserts P0 4 into the upwelling water (concentration
2xlO- 3 moles/m 3) and assumes that one-tenth of the P04 in themixed layer is removed each year (by sinking plant remains). Whatis the steady state P0 4 concentration in surface water if none ofthe plant remains falling into the deep water are oxidized (i.e.,they fall to the bottom of the column and accumulate as sediment)?
Assuming that the C/P ratio in falling particulate debris is
100 (and that no CaC03 forms) Prof. de Vendra tackles the carbon
cycle.
What are the steady state PCO l and ECO? in surface water? If the falling organic debris has tne same ~r4C as the carbon
in surface water what is the steady state ~14C for surface water?
Professor de Vendra then assumes that through its actions society raises the P0 4 content of the input water to 3xlO- 3 moles/
m3 • What would happen to the PeOof the surface water? Assume
that the PCOz content of the atmosp~ere remains at 300xl0- 6 atm.
Finally, Prof. de Vendra considers the 0 1 3 C for the surface
water in his column. He feeds in EC0 2 of o13C =at the base of
the column. He assumes that the plants Ii ving in surface waterproduce or~anic rna tter wi th a <$ 13 C value 200/00 lower than that1
for the 0 C in surface water. ¥~ also assumes that the C02 escaping f)qm surface water has a 0 C of 9 0 /00 lower than the C02 in
the 0 C in the surface water. He assumes that the CO 2 entering
the sea carries a 013C equal to that in the atmosphere over thecolumn .
Wha t i s th e 0 1 3C for surf ace wa t erEC02 ( i . e ., - 7 0 /0 0 ~in the abiotic case; in the case where the input P0 4 is 2xl0molesim 3; and in the case where the input P0 4 concentration is
3xl0- 3 moles/m 3 ?

°LAST MINUTE ADDITION (SEPTEMBER 1982)
During the last two years Richard Gammon and Joel Cline, using equipment designed at the Pacific Marine Laboratory of NOAA,
have measu'red the distribution of freon 11 and freon 12 in the
North Pacific (731, 732) and North Atlantic Oceans.
Their measurements clearly indicate the enormous potential of these tracers.
For example, Gammon found, during the TTO expedition in 1981, freons in the countercurrent flowing beneath the Gulf Stream along
the margin of the North American continent.
From the ratio offreon 11 to freon 12 he could show that the freons must have left
contact wi th the atmosphere wi thin the last five years. These
tracers have two great advantages over the tritium-helium pair.
First, they can be measured at sea.
Second, they are abundant insouthern as well as northern hemisphere waters.

443

ekapte~ 9
ICE SHEETS AND OCEAN PHOSPHATE
GLACIAL TO INTERGLACIAL CHANGES IN OCEAN CHEMISTRYINTRODUCTION
Of all the changes in ocean chemistry which may have takenplace during the 4.5 bi llion-year his tory of our planet, only
those which occurred during the past several hundred thousand
years can be documented in sufficient detail to develop a reasonably firm case his tory.
As we saw in chapter 6, the waxing andwaning of the great continental glaciers brought about changes inthe sedimentary parameters sensi ti ve to ocean chemis try.
Surprisingly enough it was, however, measurements made on polar ice
rather than on marine sediments that generated a major interest in
this subject.
Scientis ts working in Bern, Swi tzerland and in
Grenoble, France developed means of measuring the C02 content of
the air trapped in bubbles in ice and applied this method tosamples from borings made in the Greenland and Antarctic ice
caps.
As shown in figure 9-1, measurements on air extracted frombubbles in ice formed during the peak of the last glacial period
yield C02 partial pressures averaging about 200 x 10- 6 atm, while
those on pubbles from ice formed during post glacial time yield
CO 2 partial press ures ave raging abou t 290 x 10- 6 a tm. *This
discovery had a major impact on scientific thinking because of theability of CO 2 to absorb outgoing "earth" light and hence toinfluence the temperature of the earth's surface.
The results ofatmospheric modeling studies made in order to assess the climaticchange to be expected from the excess C02 produced by the burningof fossil fuels suggest that a doubling of C02 content would warm
the planetary surface by 3.0 ± 1.5°C.
A halving of C02 content
would cool it by about the same amount.
Hence the change inatmospheric C02 content between glacial and interglacial time
postulated by the Swiss and French groups would significantly
amplify the forces driving the glacial to interglacial climaticchange.
As the low C02 partial pressures of glacial time and the highC02 partial pressures of post-glacial time persisted for thousandsof years, there was in each case adequate opportuni ty for the*The ice formed during glacial times can be readily gisf~nguishedfrom that formed during post-glacial time by its 1 0 / a ratio.
q~rin% glacial time the polar regions were colder leading to lower
0/ 1 0 rat ios in the snow wh ich fe 11 at the locat ions of thesedrilling sites (see figure 9-1).

444

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9-1. Atmospheric C02 concentrations as reconstructed byNeftel, et ale from measurements of the ratio of C02 ~o air trapped in bubbles in ice cores from Camp Century,
Greenland(lower left) and Byrd Station, An tarctica (lower right).
Als 0
shown are 0 18 0 versus depth records in the same cores (upper
panels).
Approximate age scales are also given for each of theseice cores.
The nonlinearity in the age versus depth relationshipis the result of the foreshortening of each layer as the result of
the flow of the ice.
At each depth in these cores a series ofsamples were measured for their C02 to air ratio.
The range ofresul ts for these samples is shown by the vertical lines.
The
circle is the median value.
For those depths showing a large
range in results the median values are quite high.
It is thoughtthat this combination of high values and high variability is theresult of contamination with the kerosene-glycol mixture used during the drilling to keep the hole from collapsing.
Thus the resuI ts for 'those depths showing low variabili ty are the ones uponwhich the conclusions are based.
The C02 measurements were carried out at the University of Bern in Switzerland by Neftel,
Oeschger and their colleagues (622).
The change in 18 0 /16 0 ratio
wi th depth in these cores follows changes in air temperature at
the site of accumulation of snow from which the ice in these cores
was formed.
The increase in this ratio at a depth of about 1100
meters in each of these cores marks the close of glacial time.
The 18 0 /16 0 ratio measurements on samples from the Greenland ice
core were carried out by Dansgaard and his associates in Copenhagen; those on the Antarctic core were carried out by Epstein andhis colleagues at Cal Tech.
Figure445

chemistries of the ocean and atmosphere to achieve equilibrium.
Further, since the ocean contains about 60 times more carbon than
the atmosphere, the glacial to intergalcial change in atmosphericcontent must have been driven by changes in ocean chemistry.
Something must have happened wi thin the ocean which caused
themean C02 pressure in its surface waters to change!

TEMPERATURE AND SALINITY CHANGES
Two possibilities immediately come to mind.
First, the surface waters of the ocean must have been colder during glacial timeand second, the waters of the ocean must have been saltier. * It
has been possible to reconstruct the changes in surface ocean
temperature in two ~~ite different ways.
The first involves thecomparison of the 0 0 records for the shells of benthic foraminifera (microscopic animals that live on the sea floor) and ofplanktonic foraminifera (free-floating microscopic animals thatlive in the upper waters of the sea).
As shown in figure 9-2,
these records have on the average a similar full glacial to fullinterglacial amplitude.
This similarity demands that the glacial
to interglacial change in surface ocean temperature be quite
small.
The logic is as follows.
The benthic 18 0 records comefrom places on the sea floor wi th present day potential temperatures ranging down to o. 5°C:.
Since the freezing point of sea
water is about -l.Boc, the bottom waters in the sea during glacial
time could have been no more than 2. 3°C colder than they aretoday.
Since the fractionation between the oxygen in water andthat in the CaC03 formed by shells changes by about 0.22 0 /00 per°c, no more than 0.5 0 /00 of the o~~erved 1.9 0 /00 full glacial
to full interglacial change in the 0 a for benthic shells can be
attributed to temperature change.
This observed change must
rather be largely the result of the storage of 18 0 deficient waterin the continental ice sheets.
This ice storage effect would be
uniformly distributed through the sea.
The similarity between the
18 0 records for planktonic and benthic forams therefore requires
then that the glacial to interglacial temperature change insurface ocean water be about the same as that in the deep sea(i.e., <2.3°C).
The second method for establishing the temperature of surface
waters during glacial times involves a multivariant statisticalanalysis of the relative abundances of the species of various
types of ·~croorganisms (i.e., of planktonic foraminifera, ofradiolaria,
of coccolithophoridae,
and of diatoms).
In thisanalysis an equation is generated relating present-day surface
ocean temperatures to the microorganism abundances in core top
samples from throughout a major ocean basin.
When this is done it
is found that changes in surface temperature account for a major
portion of the changes in the relative abundances of the various
species from place to place in the basin. Also, temperatures backcalculated from the equation for individual locations show a mean
*Salt is left behind during evaporation hence the growth of thegreat ice sheets of glac ial time enriched the salt content ofresidual sea water.

446

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9-2. Ampli tude of the glacial to interglacial change infor planktonic (solid circles) and benthic (open circles)
forams from deep sea cores collected throughout the world ocean as
a function of sedimenta tion ra te.
The smaller values for coreswi th sedimen ta tion ra tes less than 2 cm/ 10 3 years are the resul tof a blurring of the record through biotyrbation. For cores withsedimentation rates of more than 2 cm/10 there is no significant
difference between the average amplitudes for the benthic and forthe planktonic records (197).

Ft~re
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447

deviation from the observed surface water temperataure at theselocations of only about 1°C (see figure 9-3).
Based on thesesuccesses, Imbrie and his coworkers concluded that their equations
could be reliably applied to abundance data for samples from sediments deposi ted during full glacial time.
After the long job ofanalyzing sediments from throughout the world ocean was completed
these researchers found that the difference between the averagesurface water temperature for full glacial and full interglacial
time was only about l.5°C"!
Maps showing the global distributionof these changes are shown in figure 9-4~
Imbrie and his coworkers point to the agreement between temperature differences obtained at the same point on the ocean floor from different groupsof organisms (for example by comparing those based on the abundances of various species of planktonic foraminifera wi th those
based on the abundances of various species of diatoms) as evidence
for the reliability of their approach.
As their methops yields a
result which is consistent with that obtained from the
0 record,
the conclusion that on the average the surface ocean was only 1.5± 1.5°C cooler during glacial time than it is today appears firm!
There are several methods by which the salini ty of the glacial ocean can be estimated.
The most direct is to estimate thevolume of excess (over today' s) continental ice present during
glacial times.
This is done from maps of the margins of the full
glacial ice sheets established by geologic studies.
If the ice
sheets which lay within these margins are assumed to have had anequilibrium shape* then the volume of ice can be calculated.
The resul t obtained in this way is 50 million cubic kilometers(equivalent to a sea level lowering of 150 meters). As the volumeof the sea is presently about 1330 million cubic kilometers, its
glacial volume must have been about 3.7% smaller and hence its
glacial salinity about 3.7% higher.
A similar estimate of the icevolume is obtained from the magni tude of the glacial to interglacial olS 0 change for benthic forams.
If the deep sea is assumed to have been 1.5 ± 1. 5°C colder during full glacial t~~nduring full interglacial times, then the magni tude of the 0 0
c?~nge attributable to ice growth is about 1.6%0.
If the mean
<5
0 for the excess glacial ice is taken to be -35%0 then theexcess ice volume must have corresponded to about 4.5% of thewater in today's sea (i.e., to 60 million cubic kilometers of sea
water).
Three other means of estimating the volume of excess ice
all involving the elevations of paleo-shorelines have been pursued.
The,se yield excess ice volumes in the range 25 to 50 million cubic' kilometers.
Giving some weight to all these studies,
we adopt an estimate of 45 ± 15 million cubic kilometers for thevolume of water bound up in excess continental ice during full
glacial time. This corresponds to a salinity change of 3.5 ± 1.0%(higher during glacial time).
Were these the only changes which took place in the ocean
during glacial time then the PC02 of its surface water can beeasily calculated.
The PC02 rise caused by the 1.5°C warming at*Equilibrium implies that ice is flowing toward a melting at theedges of the sheet as fast as snow is accumulating on its
interior.

448

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compared to measured surface ocean temperatures.
The map shows
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between estimated and measured temperature. This work was carriedout by John Imbrie and Nilva Kipp at Brown University (600, 603).

Figure

449

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~JO'~'~0'~50'~60~'~70~'~lO~'~90~'~'0~0"~"~O'~'2~0'~1~30'~I~'0:-'1~50~'~16O~'":":11O~'"':"::'IO~'~17~0'-:'16~O'~'50~'~I~'0':-1~30:""~120~'~1l0"~IO=0"'"'::90:::"-:8~0''''''70~'~60'~SO~'":'::.0::""~30::"" '020.::::""~IO:::'" -:0::'"-:1~O'~20::'~30"70'9-4. Maps showing the pattern of interglacial to glacial
temperature differences for the surface ocean as reconstructed
through factor analysis of the abundances of organism remains in
the underlying sediment.
Negative numbers indicate lower temperatures (in °C) during glacial time.
This work was carried out by
the CLIMAP group (610).

Figure450

the close of glacial time is almost exactly balanced by the PC02
drop caused by the 3.5% drop on se-lini ty!
(The PCO 2 of surface
ocean water rises about 20 x 10atm for a 1.5°C warming andfalls about 14 x 10- 6 atm for a 3.5% decrease in salinity). Hence
if we are to explain the ice core results, some other changes musthave occurred.

FORMATION AND DESTRUCTION OF ORGANIC MATERIALSThe most obvious possibility is that the amount of EC02 inthe sea was different during glacial time than it is today.
Were
the ECO 2 lower then the PCO 2 in surface water would have been
lower.
How might such a change have taken place?
The first
possibility which comes to mind is that one or more of the factors(mixing rate, river input rate, preservation probability of theorganic matter generated by plants, ... ) controlling the steadystate concentration of LC02 in the sea undergoes glacial tointerglacial changes.
It can be quickly shown that this is not
the explanation for the C02 to air ratio changes seen in polar ice
borings.
The reason is that while the mean replacement time ofcarbon atoms in the sea is about 200,000 years, the ice core
record (see figure 9-1) shows that the CO 2 to air ratio changed
from its glacial to its interglacial value in just a few thousand
years.
It is impossible to envision a scheme by whic~hanges in
the parameters controlling the steady state LC02 content of sea
water could produce such an abrupt change!
Instead, we must look to pulses of formation and destructionof temporary carbon reservoirs.
Forests, soils, and shelf sediments come to mind.
What evidence would such transfers leave?
In
the case of reservoirs of organic carbon, the isotopes of carbon
come to mind.
As we have seen, the heavy isoto~e, 1 3C is discriminated against during photosynthesis.
The 1 C/ 12C ratio inplants is about 20 0 /00 lower than in the inorganic carbon
utilized by the plant.
Thus just as any prowth or depletion ofcontinental ice leaves its mark on the 180 / 6 0 ratio in sea water,
any growth or de~letion of reservoirs of organic carbon will leave
its mark on the
3C/12 C ratio in oceanic EC02.
The production and destruction of reservoirs containing
CaC03 would of course also change the ocean's EC02. They would in
addition change the ocean's alkalinity.
They would not, however,significantly change the 13 C/12 C ratio in ocean-atmosphere carbon.
As the fr~ctionation between the carbon in CaC0 3 and the dissolved
ca rbon in· the sea is only abou t 1 0 /00 (c ompared to 20 0 /00 fororganic material), the 13 C record speaks only to changes in thesizes of organic carbon reservoirs.
We will have to look elsewhere for evidence regarding changes in the size of the sedimentary CaC03 reservoirs.
The best place to look for the changes generated by theformation and destruction of reservoirs of organic matter would bein the shells of benthic foraminifera.
Their carbon comes fromthe deep sea reservoir which contains about 90% of the carbon in
the ocean-atmosphere system.
They live in an environment nearly
free of temperature changes.
They do not migrate through the water column as do planktonic foraminifera.
Hence changes in their451

1 3 e/ 12 e ra tio should mos tnearly reflect those for the ave rage
carbon in the ocean-atmosphere system.
In chapter 6 we showed that a glacial to interglacial change
in the 0 1 3e of benthic forams has been found (see figure 6-10).
Forams Ii vin~ in glacial periods were observed by Shackleton tohave lower I e/ 12e ratios than those living during interglacial
periods.
Table 9-1 summarizes the resul ts available at the time
tq~s book was written.
All show a lower glacial than interglacialo e value. The average difference is 0.7 0 /00.
When the data in figure 6-10 are examined it becomes evident
that the glacial to interglacial shift for the two species analyzed are not the same (see figure 9-5).
Hence a word of caution
must be given regarding the acceptance of the 0.7 0 /00 average.
Benthic forams may have m~re surprises in store for us!
If we accept the 0 1 e resul ts at face value the reservoirs
containing organic carbon must have been smaller during glacial
time.
This would lead to higher eo 2 contents of the ocean andatmosphere during glacial time!
The magni tude of the difference
can be estimated from the magnitude of the o13 C change observed inbenthi~ forams between glacial and interglacial time.
Takin~ themean 0 3 e value for the organic carbon to be -23 ± 3 0 /00 ,itturns out that 3.0 ± 0.4% of the carbon in the glacial ocean musthave been converted to organic carbon at the end of glacial time.
Were this the only chemical change to have occurred in the sea,
the C02 partial pressure in the atmosphere must have fallen about
150 x 10- 6 atmospheres at the end of glacial time!
As the iceboring data suggests instead that it rose by 100 x 10- 6 atmospheres at the close of glacial time we must probe further.
Before going on, we must mention where the excess organic
material formed during early post-f,lacial time might be stored.
Shackleton, who discovered the 0 1 e change in benthic forams,
proposed that this carbon is locked up in fores ts and soils.
He
points out that during glacial time the north temperate fares ts
which now store about half the carbon in the living biosphere werelargely non-existent. The land they now occupy was either coveredby ice or by treeless tundra.
He also points out that the tropical forests appear to have covered a smaller area during glacial
time because of generally drier conditions in the tropics. As canbe seen from the data in table 9-2, glacial to interglacialchanges in forest cover alone cannot account for the trfnsfers ofcarbon needed to explain the glacial to interglacial 0 3C change
in benthic forams.
Shackleton realized this and proposed thatchanges in" the amount of organic carbon stored in soils must also
have occurred.
As can be seen from table 9-2, there is enoughsoil carbon to do the remaining job.
Another possibility is that deposition of organic richsediment on the world's shel ves occurred during the post-glacial
sea level rise.
During the peak of glaciation, the sea stood*Terrestrial organic mat ter hf~ a 013C value averaging -26 0 /00.
Marine organic matter has a 0 C value averaging -20 0 /00.
As wehave no way of determining the relative contributions of the two
types, we take a value half way in between and an uncertainty
large enough to bridge the entire range.

452

Table 9-1. Summary of 13 C data on benthic foraminifera.
These results were all obtainedby Nick Shackelton of Cambridge University in England (328).
Core
No.

Lat ..

Long.

Deptn
km

13C
(ECOz)*

Species

0/00

Glacial
13 C<5Holocene
o13 C

0/00

0/00

/1 <5
CG-I
0/00

13

Atlantic~c..nV29-17944.0 0 N

24.5°W

3·3+0.9

UVig. per.

-1.1+

-0.2+

+0.9+ )

Meteor12392

25.2°N

16.8°W

2.6+0.9

Uvig. per.
Plan. woel.

-1.0
+0.3

+0.2
+1.0

+1.2
+0.7

V22-174
RC12-267

10.1°S
3H.7°S

UVig. aub.
UVig. per.

0.1
-0.9

+0.5
-0.1

+0.4
+0.8

12.8°w
25.8°E

2.6
4.1

+0.8
+0.3

I
0.8

CA:>

Pacific
V19-293.6°8

83.9°W

3.2. . 0.0

UVig. probe

-1.45

-0.75

+0.7 }

Y71-6-12MG

16.3°S

77.5°W

2.7. . 0.0

Uvig. per.

-1.3

-0.8

+0.5

-0.9

-0.1

+0.8

MEAN

+0.7

0.6

Antarctic
RCll-120

43.5°S

79.9°E

3.2+0.4

Uvig. per.

+Across termination II (125,000 years ago) rather than termination I(11,000 years ago).

*For the carbon dissolved in the sea water at the depth and location the core was taken.

al8 0

al3 c

5.0

0
1.2%.

f<<:> ~
d~
0

I
00 I
0 0

~ O.7%o~

0

2

3

~
%1
~~

-E
-

0§

0

5

W
0

6

,

0~

0

0

0

<:e

0

0

0

o0 0
o 0 0

~

0
0

~~

*"* *

***<>

~

7

*

8

••
*

o 00

I0

J:
f-

a..

I

•1*

0

* ?'.
~

.* *1

*#~}
*

o 1.4°/00°;f.'~rJ0P

TERM~0 §II

0~Figure 9-5~ Plots of 013 C and 0 18 0 versus depth in METEOR core12392 from the eastern North Atlantic.
Analysis carried out on
Uvigerina peregrina are designated by circles.
Those carried out
on Panulina woellerstorfi are designated by stars.
To make thePanulina woellerstorfi results compatible with those for Uvigerina
peregrina 1.0 % 0 was subtracted from the Panulina woellerstorfi
o13 C results.
The last two transitions from glacial to interglacial time are designated as term:\nation I and II. The glacial tointerglacial change for the 0 3 C of
Uvigerina peregrina is
1.2 0 /00, that for Panulina woellerstorfi is 0.7 % 0.
The reason for this difference is not known.
These measurements were madeby Shackleton at Cambridge University (328).

454

Table 9-2. Today's reservoir sizes and the magnitude of glacial tointerglacial carbon transfers.

Reservoir

Mass Carbon
10 16 moles

Ocean & Atm.

310

Forests

6

Soil

~20

to Interglacial
Organic C Storage Change

9

Glacial to Inter~lacial
CaCa 3 Compensation

9~Glacial455

about 140 meters below its present level leaving the shelves highand dry.
Thus during glacials it is likely that material was
eroded from the shelves.
On the other hand, as the sea rose andreflooded the shel ves, net deposi tion would be expected.
This
could contribu te to the glacial to interglacial cycle in organiccarbon storage and hence to the o13 C change observed in benthic
forams.
Whether terrestrial organic materials (forests and soils)
or marine organic material (shelf sediment) dominated the change
in the glacial to interglacial inventory of organic carbon will be
important to the arguments which follow.
Thus we will have more
to say about this matter.

CHANGES IN CaC0 3 STORAGE
What about the changes in the C~CD3 reservoirs? The removal
of 3% of the carbon in the ocean-atmosphere system at the end of
glacial time to form organic material would have produced an increase in CD 3= ion concentra tion.
In the extreme (i. e., .were
there no compensation due to excess CaCD3 accumulation during theperiod of storage of this organic carbon) the CD3= ion content
would have risen about 15 x 10- 6 ~m/kg.
As can be seen from figures 2-10, 11 and 12 such an increase would have caused almost theentire deep sea to become supersaturated with respect to calcite
and would have caused the aragoni te lysocline to deepen by morethan one kilometer!
Such a change would of course throw theocean's CaC03 budget badly out of balance, leading to an excess of
CaC03 accumulation over input from rivers of carbon destined forthis fate.
This excess accumulation would gradually bring down
the
C03=
ion
concentration and eventually
restore balance
between input and loss.
The loss of EC02 to organic matter wouldin this way be "compensated" by an equivalent loss of LC02 to
CaC03. For each mole of carbon removed as organic material about
a mole would be removed as rraCD3.
When the CaC~03 compensation was completed the ocean would
have lost an additional 3% of its carbon. It would also have lost
about one mole of alkalinity for each mole of EC02. As precipitation of CaCO 3 raises the CO 2 pressure in sea water, the C:aC:O 3
compensation would have countered much of the CO 2 pressure reduction caused by the loss of LCO 2 to organic carbon.
Thus thiscompensation would undo most of the difficulty created by theobserved o13C change in benthic foraminifera!
Befo~e looking to other causes for the C02 pressure increase
observed in ice cores at the end of glacial time, there are some
aspects
of
this
CaC03
compensation process
which we must
consider. The first is how long it would take. A rough estimate
can be quickly obtained as follows.
The replacement time forcarbon in the sea is about 200,000 years.
However, as the ratioof EC02 ion to C03= in deep ocean water is about 30, theresponse time of carbonate ion will be about 30 times smaller(hence about 7000 years).
A more easily understood way to estimate the compensation
time is to consider how much excess CaC03 has to be precipitatedand compare this wi th the ra te of productiq~ of CaCO 3 by marineorganisms. The ocean contains about 3 x 10
moles of ECO 2.
As

456

the compensation Pl~cess required the deposition of 3% of thiscarbon, then 9 x 10
moles of excess CaC03 had to be formed.
Theaverage production rate of calcite is 1.0 g/cm 2 10 3 yrs, or 100
moles/m 2 10 3 Irs.
Over the entire ocean this comes to 3.5 x 10 16
moles per 10 years.
If one-half of this production (i. e. abou t
1.8 x 10 16 moles/10 3yrs) went into the accumulation of excess*
CaC03 then the time constant for the compensation would be about
5000 years.
We refer to it as a time constant rather than as atime because as C03= ion content of the sea fell the area
covered by supersaturated water would correspondingly diminish.
As for all the feedback recoveries discussed in chapter 6, fullrecovery of the CaC0 3 balance would be approached asymptotically.
The time estimates gi ven above are the so-called e-folding times(i.e., the time required for the departure from equilibrium to be
reduced by a factor of lie).
The half recovery time would be In2(or 0.693) times the e-folding time.
(Hence the half time for
CaCO 3 compensation would be 0.693 x 5000 or 3500 years).
In anycase, the rough estimates made here suggest that the compensation
would persist well into the Holocene.
Possibly it still goes ontoday!

EVIDENCE FOR AN EARLY POST-GLACIAL LYSOCLINE CHANGE
Can we see evidence for this compensation process in thesediments?
The answe r is yes.
Spikes of aragoni te preserva tionfound at the appropriate level in sediments from the Caribbean Sea
and from the eastern basin of the Atlantic are shown in figure~-6.
These cores were all obtained beneath the present-day
lysocline for aragonite.
The fact that aragonite is absent fromthese cores during glacial time suggests that the lysocline foraragoni te resided at depths shallower than the collection depths
for these cores during glacial time as well as during Holocene
time.
The preservation spike then denotes a period several
thousands years in duration during which the aragonite lysocline
was located at a greater depth.
The timing of this drop iscentered about the glacial to interglacial transition.
In thecase of the eastern Atlantic study a series of cores were taken
down the continental rise over the depth interval 0.5 to 3.7 km.
As shown in figure 9-7 by comparing the results on all these cores
it is possible to roughly determine how far the lysocline foraragoni te dropped during the preserva tion event.
As can be seenfrom this. figure the drop was qui te large (i. e., about 2 kilometers).
Hence we have a dramatic confirmation that the removalof IC0 2 from the ocean in the form of organic matter at the closeof glacial time caused a response in the aragoni te lysoclines.
However, as aragonite is quite rare in sediments it cannot be thesought after storage phase for CaC03.
If we are to rigorously demonstrate the CaC03 compensation
phenomenon it must be through an evaluation of the extent ofexcess calcite storage.
This is not so easily done.
In order to*It is an excess because this deposition would have to be over andabove that to balance the input from rivers of carbon destined to
become CaC03.

457

al80
0···.....-...., ....- ....... 0 .....-.--.........

o

2

r

r:==_-~...

-..)("'

...

,,0
Cl

r

l>
0
l>

.....-.-

E 2........

E........,

r

'-'

I

:r:

....a..
W

I-

3

3

J>

0...
W

(J)

0

tD

0

",

z

~

t

4

4

z

2

~

rrt
~

Cl

r

5

METEOR

V12-122512310-4

0

10 30
PTEROPODS (%)
1

2

5

3
I
PTEROPODS (°/0)

0

4000

3
8000

NO. PTEROPODS
g COARSE FRAC.

Figure 9-6. Pteropod abundance (solid circles) as a function of
depth in a core from the Caribbean Sea (501) and in two cores fromthe continental margin off Af rica (503,509).
All three records
show preservation peaks centered in early post-glacial time.
Bothsites lie below the present day aragoni te lysocline.
The timescale for core V12-122 is based on 18 0/ 16 0 measurements (see x's).
Those for the other cores are based on faunal evidence.

458

Figure 9-7. Changes in the lysocline depth fo~ pte~opods as reconstructed for a se~ies of cores from off the coast of Africa.
Thesuggestion is that the depth of this horizon increased by about 2
kilometers at the end of glacial time and then decreased by about
3 kilometers during mid-Holocene time. The entries are for depths
where cores were obtained.
The designations at each of these
depths are for the relative abundances of pteropods in late glacial
sediment, in early Holocene sediment and in late Holocene sediment. This reconstruction was carried out by Berger of the Scripps
Institution for Oceanography (511).

HOLOCENE TRANS GLACIAL-E-VE-;Y

-

oX

-;A~

VERY RARE

:I:
l-

'\

COMMONCOMMONI COMMON COMMONPTEROPODI
Ir - - - LYSOCLINEI

a.
w

ABSENT

I COMMONa::

ABSENTIABUNDANTIABSENT

ABSENTIABUNOANriVERY RARE

Cl

w

l-

~

3

ABSENT
ABSENT

I VERY RARE

\c\8U~A~TI
ABSENT

VERY RARE
ABSENT

4
EASTERN NORTH ATLANTICOFF COAST OF AFR leA

Table 9-3. Areas covered by various sedimentary domains (inunits of 10 14 m2 ).

Zone

Atlantic

IndianPacific

Total

Above Lysocline

0.71

0.32

0.49

1.52

Trans. Zone

0.21

0.14

0.14

0.49

Red Clay Zone

0.14

0.30

1.15

1.59

Total

1.06

0.76

1.78

3.60

459

put thislayer ofquest into context let us first consider how thick a
calcite we seek.
A rough idea can be obtained by
diy~dlng the total amount of CaG0 3 required for compensation (-9 x
10 2 moles) by the total area of the deep sea floor (-3 x 10 18cm ).
This yields about 3 x 10- 2 moles CaG0 3/cm 2 or 3 grams of
CaGa 3/cm 2.
Since a cubic centimeter of calcite rich deep sea
sediment contains close to one gram of mineral matter if uniformly
spread over the entire deep sea floor, this CaC03 would produce alayer about 3 em thick.
Because the layer would not be formed
everywhere the actual thickness would be somewhat larger.
For the purposes of our search let us di vide the sea floor
into three regions: that above the present-day lysocline for
calci te, that lying in the transi tion zone beneath the lysocline
and that beneath the transi tion zone.
The areas of these zones
are given in table 9-3.
The zone above the lysocline for calcite
will not show any evidence for the compensation phenomenon.
Thereason is that the phenomenon is one of preservation rather thanof deposition.
All the calcite raining above the lysocline ispreserved even in the absence of such an event!
Thus a temporary
deepening of the lysocline will not alter the accumulation rate of
calcite above the lysocline.
Next, let us turn our attention to the sediments nearly freeof calcite which lie beneath the transition zone.
Here we might
expect to observe a calci te preservation spike similar to thearagonite preservation spike discussed above.
The situation is,
however, quite different because of low accumulation rates in thisdomain.
The aragoni te preserva tion spike is found in sediments
accumulating at the rate of several centimeters per thousand years(most of the bulk in these sediments is calcite and aluminosilicates) .
In such sediments burrowing organisms stir only the
sediment laid down in the las t seve ral thousand' years.
Hence,
while s lightly blurring the preservation spike, the bioturbation
does not obliterate it. Beneath the transition zone the depositon
rates are only a few tenths of a centimeter per thousand years.
In this domain
the entire column of Holocene sediment is being
actively churned into the underlying glacial sediment!
Let us imagine what would happen if a 6 cm thick layer ofcalcite were deposited during Holocene time on the top of glacial
red clay.
Through the ac tion of the worms Ii ving wi thin thesediment this calcite would be thoroughly mixed with the red clay
deposited during late glacial time and with the red clay deposited
during Holocene time.
Thus, unlike the aragonite preservation
event which appears as a discrete layer within the sediment, thecalcite preservation event would be largely homogenized throughout
the upper eight or so centimeters of sediment.
It would gi ve
every appearance that calcite accumulation had commenced at theclose of glacial time and continued right up to the present.
Indeed, how could these two scenarios be distinguished? In both,the fauna and flora in the sedimen t would be Holocene in age.
Further, as continuous deposition over the Holocene represents a
simpler hypothesis, its adoption would be favored by the usual"rules" of science.
There is, fortunately, one way in which the two hypotheses
might be distinguished.
Were the calcite to have been deposited460

continuously from early Holocene time until now its average radiocarbon age would be about 5000 years (the transition between glacial and Holocene time occurred about 11,000 years ago).
On theother hand, if it were depos i ted mainly during early Holocene
time, its mean age should be considerably older (8000 to 10,000 or
more years). So we should look at sediments from beneath the transition zone to see whether they have an 8 or so centimeter thickcarbonate bearing layer at their top.
If such a layer is foundits calcite should be subjected to radiocarbon dating.
Such a study was carried out on sediments from beneath thelysocline on the west flank of the East Pacific Rise.
In thisarea of the ocean the sediments above the lysocline are very richin calcite (90-96%) (except at the ridge crest where they are
diluted wi th the oxides of Fe and Mn).
As shown in figure 9-8,
prominent solution effects commence at a depth of about 3.95 lan.
Sediments from below this water depth have a layer of CaCa 3-richsediment on top of the underlying red clay (see figure 9-9).
These layers are a few tens of centimeters thick, and yieldradiocarbon ages in the range 9,000 to 17,000 years (see figure
9-9 and table 9-4).
These layers are considerably thicker thanthe expected global average and have radiocarbon ages several
thousand years older than expected.
This si tuation is further
complicated by the finding that core tops from above the lysocline
have ages in the range 3,000 to 16,000 years.
When considered to~ether wi th the 23 °Th based sedimentation
rate as low as 0.4 cm/10 years for cores from above the lysocline
cores (see table 9-4) another explanation for both the high 14C
ages on both super - and sUblysocline core tops comes to mind.
If
the bioturbation depth is 8 em, then the mean 14C age of the mixed
layer in superlysocline cores would be about 20,000 years.
So then what is the origin of the sUbly socline calci te richcore top material? The integrated mass of calcite in these layers
comes out to be 10 to 25 g/cm 2 .
At a rain rate of 0.4 g/cm 2 10 3
yrs, the accumulation of this material would take 25 to 60 thousand years.
Furthermore, as this sublysocline core top material
appears to have been SUbjected to calcite solution the time foraccumulation must have been even longer.
Thus we are likely
looking at a change in lysocline depth on the time scale of muchgreater than that for the Holocene.
Our calcite preservation
event may also be there,· but its message is hidden by the complexity of the situation.
Unfortunately, we cannot use the resultsfrom sediments in this area of the ocean ei ther in support of orin opposition to an early post-glacial calcite preservation
event.
The third sedimentary domain is the transition zone in which
partial calcite solution occurs.
In this zone the calcitepreservation event should appear as both a maximum in accumulation
rate and a minimum in extent of dissolution.
An example of the
GaGO 3 versus depth record for a core taken from such an area isshown in figure 6-e, along with an 180 record for the same core.
While the major feature of this record is an indication (throughlower CaCO 3 contents) that dissolution was more intense during
interglacials than during glacials, there is also a marked lag
between the CaC03 and 8 18 0 changes.
This lag is most easily seen461

Figure 9-8. CaC03 content (% by weight) at depths of 10 and 30 cen-timeters for cores from the west flank of the East Pacific Rise
plotted against the water depth at which the core was collected.
For the samples taken at 30 crn a sharp transition occurs at a water depth of 3.95 km. This research was carried out at the LamontDoherty Geological Observatory (4e7).

462

CaC0 3 (O/o)

0050r----..

E

Co)

20

100

V19-5417°S 1140 W,'
2965M'

3500

°

0

5500

...---.0lIl..--.....
V19-6317°S IISOW
3445M. 980010,70014.700/5.60014.500

:I:
~

~ 40

o

1.2cm/103y

SED.RATE.

SED. RATE
0.4 cml10 3 yrs(BASED ON(BASED ON

230Th)

230 Th )60 '-'_........ ~._ ...

0..---......- ......
12,90014,500

E

~
:I:
I-

20

17,600

a.

w 40

o

60 ....._-..110-_.-..

Figure 9-9.

Plots of CaC03 (in weight %) against depth in the sediment column for cores taken on the west flank of the East Pacific
Rise.
Radiocarbon ages on the CaC03 are also given. V19-54 andV19-63 are from above the lysocline. The others are from beneaththe lysocline. These measurements were made at the Lamont-Doherty
Geological Observatory (487, 512).

463

2..--......- - - -......----.,......,
2

V12- 122--;f!.

--I ••G. RU8ER........ .... . .. ....

•

o

-I
U

••

'.I'

ttl
(,()

•

A.

• • • • tr'Y.

1.1

...\

--.

•

.,~_.

_.

V12-122

~~

-~--~

~.

•• •
•

•

G.SACCULIFER

o

1

•

o
COLD

••

......

I, . ••

~

WARM

2.0

• •

•

-. ,...
. . -.-~1.95--.:~)o• ••(;, • •1.8•

1.5 I-••• ERDC 92+128
G. SACCULlFER
-2.0

WARM

G. SACCULIFER-AND G.RUBER• . -• ••. •~1.7~•••• • • •

1.5

•
••
•

RC17-177 I-+--1.6~

•

•

• •

-.

-

•

-1.5

8 18 0(0/00)

1.0

-1.0

COLD

I

I

I

-2.0

- 1.5

-1.0

WARM

al8 0 (%0)

COLD.

Figure 9-10.
Plots of c 1 3 C ve rsus C 1 8 0 for planktonic foraminifera from Caribbean Sea core V12-122 (327) and from western equatorial Pacific core EROC 98 (329), EROC 128 (329) and RC 17-177(personal communication, Shackleton).
Averages for samples fromglaciatime (high IS 18 0 ) and for samples from interglacial time18
(low IS 0) are shown.

464

Table 9-4. Core top radiocarbon ages for sediments on the west flank of the East Pacific
Rise.
These measurements were made at the Lamont-Doherty Geological Observatory (487,
512).
Core
No.

Lat.

Long.

Water
Depth
km,'

Depth
In corecm

CaCO 3

%

14

C
Age
yrs

Sed.*
Rate
cm/10 3 yrs

Above Lysocline
V19-53

1 'loS

114°W

3.06

2-6

V19-54

17°3

114°W

2.96

0-5

V19-61

17°S

117°W

3.42

V19-63

17°3

121°W

3.45

3,300

-

61**

3,500

1.2

0-5

90

16,600

0.4

6-10

92

9,800

0.4

Below Lysocline

~

CJ")

C..T1

V19-66

16°3

128°W

4.10

2-8

61

12,900-

V19-67

16°S

130"W

3.95

1-4

78

10,700-

V19-69

16°s

134°W

3.97

2-6

83

17,700-RC13-74

18°s

146°W

3.96

0-5

70

12,600-RC13-80

19°5

127°W

4.03

0-6

14,800-RC13-81

19°5

124°w

3.75

0-2

14,800-

V18-302

13°S

141°W

4.13

5-9

12,800-

RC15-25

70 S

137°W

4.48

0-2

9,400

-

V18-295

16°3

150 0 W

4.22

3-8

16,400-65*Based on 230 Th dating.
**Dilution by Fe and Mn oxides rather than depletion by calcite solution.

for the three warm peaks of the last interglacial period and for
the Holocene. In each case the CaC03 change occurs well after the
18 0 change.
It is possible that we see in this core the combination of two phenomena. First, the steady state lysocline level inthis area of the equatorial Pacific was deeper during glacial thanduring interglacial times.
Second, a calcite preservation event
temporarily counters the change in the steady state lysoclinelevel and in so doing generates what appears to be a delay inresponse. Again,while a tantalizing hint that a calcite compenastion event occurred, verification and quantification await further
studies. And so it stands; we have hints but no solid evidence.

QHANGES IN PHOSPHATE CDNCENTRATIONAs the title of this chapter implies, phosphate must have apart to play in this story. To understand why this is the case weneed to understand the role of phosphate in controlling the C02
partial pressure in today's surface ocean. As shown in table 9-5,
if deep sea water were brought to the surface and merely warmed
without any biological activity (i.e., heated isochemicallY)i thenits C02 partial pressure would be about three times higher thanthat in the present day atmosphere. The removal of ~C02 by plants
draws down the CO 2 pressure in the wa ter.
The extent of thisdrawdown depends on the P04/~C02 ratio in deep water. The greater
this ratio the greater the reduction in C02 partial pressure due
to plant growth!
Thus a glacial to interglacial variation in theratio of P0 4 to ~C02 in the sea would lead to a corresponding
variation in the partial pressure of CO 2 in surface ocean waterand in the atmosphere.
How would such a change be recorded in deep sea sediments?
We saw in chapter 6 that the distributions of <S 3C and P0 4 wereclosely correlated in today's ocean (see figure 6-11).
This isbecause the main source of 13 C/12 C variations from place to place
within the sea is the photosynthesis-respiration cycle.
The
difference in the 13 C/12 C ratio between phosphate-free surface
water (i.e., PFSW) and mean deep water (i.e., MDW) is given by the
relationship:
13
<S

In
b

13

CPFSW -

13
<S

CMDW

today·'. s
ocean
the
0 /00;
is-20
Cphoto,9-1

photosynthetic
fractionation factor,
the carbon to phosphorus ratio in*This might be thought of as the "Strangelove" effect.
Were lifein the ocean to be suddenly destroyed, then the ocean's surface
chemistry would change on the time scale of a few decades to thatshown for isochemical warming in table 9-5.
Such a disasterappears to have occurred at the end of Cretaceous time due to an
impact of an asteroid or comet.
Some of the effects describedhere should be sought in the records for t~at time (i.e., a C021
induced warming, a shift in surface water 0 C values, ... ).

466

Table 9-5. Interglacial to glacial changes in ocean chemistry.

P
atm

6

S

°c

0/00

Alk
lJ,eq/kg

l:C02
lJ,m/kg

C0 3=
lJ,m/kgP04
lJ.ffi/kg

NO 3
lJ,m/kg

2.2

33

PCO 2
10- 6 atm

Average Deep Water.....::l

400

1

34.7

2365

2250

86

-

~

H

Brought to the Surface and Isochemically Heatedo~.....::l
CJ

a::

1

22

34.7

!iJ

2365

2250

101

2.2

33

1043

0.0

0

266

0.0

0

311

3. 2

48

After Loss of P0 4 (and N03) to Plants

8

Z

H

1

22

34.7

23ge

~

2019

269

Arter CaC03 Loss

0')

'-oJ

1

22

34.7

2282

1961

226

Average Deep Water

390

1

35.9

2577

2466

89

Brought to Surface and Isochemically Heated.....::ll
~

~I

1

20

35·9

2577
After Loss

1

20

35.9

2625

2466

103

3.2

48

1151

0.0

0

194

0.0

0

233

of' PO 4 (and NO 3) to Plants

2130

354

After CaC03 Loss

1

20

35.9

24~7

2046

289

marineplants,C/P)org,
is about105;
the mean phosphateconcentration in deep water, [P0 4 JMDW, is 2.25 llm/kg; and themean dissolved inorganic carbon concentration in deep water,[EC02JMDW, is 2250 llm/kg.
Hence:
013CPFSW - 013 CMDW = 20 x 105 x 2.25/2250 = 2.1

0/ 00

If we assume that the fractionation during photosynthesis andelemental composi tion of marine plants was the same during~eak ~lacial time as it is today, then the difference between the
3C/ 1 C ratio in glacial phosphate free surface water and glacial
mean deep water provides a measure of the ratio of the phosphateconcentration to dissolved inorganic carbon concentration in the
deep water of glacial time. This
13 C/12 C ratio difference can bereconstructed by studies of planktonic forams and benthic forams
which lived during peak glacial time.
The temperature dependence of the carbon isotope fractionation between CaC03 and the dissol ved inorganic carbon in sea
water is quite small.
Also, as we have seen, the temperaturechanges in both surface water and deep water between glacial andinterglacial times were on the ave rage qui te small.
Hence thedifference between the 13C/12C ratio in the growing shells and in
the dissolved inorganic carbon in the water in which they grewshould not have been significantly different during glacial tq~nduring interglacial time.
Hence, if we see a change in the 0 C
value for a gi ven species of ei ther benthic or planktonic foram
between glacial and interglacial time, there is a good chance thatit reflects a change in the 013C for the dissolved inorganic carbon in that part of the ocean.
A problem to be faced in this connection is that each species
of foraminifera shows a somewhat different fractionation factor.
Geochemists term these differences "vi tal effects" (a name whichhides the fact that we don't understand them).
Because of th sevital effe~ts we cannot simply estimate the difference in the 01 3C
between glacial phosphate-free surface water and ~lacial mean deep
water directly from the difference between the 0 1 C values for theshells of planktonic organisms grown in phosphate-free surface
water and for the mean of shells of benthic organisms grown invarious places in the deep sea.
Rather, we estimate thisdifference in a less direct way from the following equation:
the(013CPFSW,- 013 CMDW )G = (013CPFSW - 013 CMDW )
+
( 13
13C)
13
13
today
o CG-oI plank - (0 CG - 0 CI)benthic

9-2

The term on the left hand side of this equation is the answer weseek.
The first of the three terms on the right is the measured
o13C difference between phosphorus-free surface water and mean
deep water in today's ocean.
The second term on the right is theglacial to interglacial change for a given species of planktonic
foraminifera as measured in individual deep sea cores.
The third
term on the right is the glacial to interglacial change for agiven species of benthic foraminifera as measured in individual
deep sea cores.
If the vi tal effect for a gi ven species isassumed to remain constant from glacial to interglacial time, its

468

influence is eliminated by this procedure.
We have already shown that the glacial to interglacial IS 1 3 C
difference for benthic specirs collected from throughout the deepsea averages 0.7 0 /00 (the 1 C/ 12 C ratio is lower during glacial
than during interglacial time).
It remains only to establish thisdifference for phosphate-free surface water.
In figure 9-10 are
shown plots of IS 13 C against IS 18 0 for different species of planktonic foraminifera in cores from the temperate tropical ocean(hence from beneath phosphate-free surface water). Samples having
high 0 18 0 values grew during times of large ice cover (glacials)
and those wi th low 0 18 0 values grew during times of small ice
cover (interglacials). It can be seen from these plots that while
the scatter of individual points is large, the difference between
the average 013C for the glacial and interglacial samples is quite
small.
These differences as summarized in table 9- 6 average only
0.1 0 /00.
Hence, as do benthics, the planktonics grown during
glacial time have a lower 13C/12C ratio than do those grown during
interglacial time. However, in the case of planktonics, the difference is far smaller and might even be said to be not significantly different from ze1o.
It is because the IS 3 C difference for planktonics is so small
and the scatter of individual measurements so large that the modeof presentation in figure 9-10 was selected. While this procedureallows us to pick out a small signal in the presence of considerable noise, the question as to the cause of the noise remains.
Were it just a matter of random variation this noise could beovercome by averaging large numbers of analyses as has been done
here.
The problem is that the factors generating the sample tosample variation might produce a bias in the glacial to interglacial difference (i. e., the difference obtained from the analysis of forams might not be the same as the difference for the
carbon dissolved in the water).
While no one has yet figured out
why 013 C analyses on planktonic forams show greater scatter thanthose on benthic forams, there is one clue.
When different size
fractions of the same species are analyzed, the smaller shells
show lower ol3 C values than the larger shells (see figure 9-11 foran example).
As these organisms add chambers to their shells as
they age, the smaller shells represent organisms that died at a"young" age, and the larger shells represent organisms that diedafter a full life. Hence the 013 C values for the calcite deposited by these species increase as they grow older.
This could be
ei ther the. result of a change in vi tal effect wi th rna turi ty or ashallowing 'of their depth habitat with maturity (i.e., less shellmaterial ~rown in the lower o13 C thermocline water and more in thehigher 0 I C
mixed layer water).
The latter explanat~qn is unlikely as the 18 0 change is in the same sense as the
C change(see figure 9-11). Were these changes produced by vertical migration the 13 C and 18 0 changes should be in the opposite sense. **The 13C/ 12C rat io in EC02 decreases wi th depth.
As the t ernpfr'i.t,pre dependence ofthe reo 2 - CaC03 fractionation is small ~o f~sO should theC/ l; ratio in shell
rna terial grown at var ious de~~hs.
The 1 0/
rat io rema ins nearly the s arne wi th
depth.
However, the H20 0/160
frl'\ctioEPation increases by about 0.2 % 0 foreach degree the water cools.
Hence the~Ll? ratio should increase with the depth
at which the CaC03 is deposited while the 3C/ 2C ratio should decrease.

°469

Table 9-6.
Core
No.

13 C

Ref.
No.

results on planktonic foraminifera.
Lat.

Long. Depthkm'

o 13 C++Species( EC 02)
0/00

Glal~aloC
0/00

Holol~ne
o
C

bo

13

CG+r

0/ 00

0/00

Caribbean
V12-122

(327)

17°N

74°W

2.8

-2.7

ruber.

0.7

0.7

0.0

V12-122

(327)

17°N

74°w

2.8

-2.7

sacca

1.3

1.1

-0.2

Atlantic
V22-174(328)

10 0 S

13°W

2.6

-2.7

ruber.

1.0

1.4

+0.4

V22-174(328)

10 0 S

13°W

2.6

-2.7

sacca

1.8

2.0

+0.2

~

o'"'

Pacific
ERDC-92

(329)

2°S

157°E

1.6

-2.7

sacca

1.8

1.95

+0.15

ERDC-128 (329)

OOS

161°E

3.7

-2.7

obliq.

1.2

1.1

-0.1

RC17-177

2°N

159°E

2.6

-2.7

sacca
+ rub.

1.6

1.7

+0.1

MEAN

+0.1

t

Shackleton (pel~onal communication)++In 1973 the 0 C for warm surface ocean water averaged 1.9 0 /00.
The recent influx of forest, soil and fossil fuel C02 has reduced the ratio by about 0.8 0 /00.
Thus prior to 1850 it was likely close to 2.7 0 /00.
t

2355-425fL••••..........
0~

"-"

U

.-•

355-425 fL

212-250 fL

0

0

rt')

-GO

G. RU BER
CORE RC 14-36

e.
-I

-2
-3

-2

212- 2 50 fL

SLOPE OFEQU I L.

G. BULLOIDES

TRENDCORE E45-70

a

-I

2

3

al80 (%0)
Plots of 013 C versus 0 18 0 for different size fractions of the species G. ruber (upper left) and G. bulloides (lower
right) from core top samples from two si tes i n the Indiaq Ocean
cores.
In both sets, the smaller shells have both lower 0 8 0 andlower 013 C values than their larger counterparts. The trend withsize does not correspond with that expected if the organi mschanged their depth habi tat wi th age.
In this case the 01 3C
18
should decrease with increasing 0 0 following the trend shown by
the large arrow.
The results were obtained by Matthews and hiscolleagues at Brown University (330).

Figure 9-11.

471

In any case, we have no way to know how glacial planktonics differed from their interglacial brothers, ei ther wi th regard tovital effect or depth habitat. Thus we must fall back on the hopethat by analyzing enough different species at enough different locations on the sea floor any bias introduced in this way is averaged out.
If we accept the glacial to interglacial differences obtainedfor planktonic and for benthic forams as meaningful estimates of
the changes in the 013 C value for phosphate-free surface water andfor mean deep water, then the difference in 013C between thesewaters in glacial times becomes:(0 1 3 CpFSW - 0 1 3 CMDW ) G = 2.1 - 0.1 + 0.7=

2.7 0 /00

These changes are shown diagramatically in figure 9-12.
If we
further assume that the magni tude of the photosynthetic isotope
fractionation factor and of the carbon to phosphorus ratio inmarine plants during glacial time was the same as during interglacial times, then the mean phosphate content of deep ocean water
during glacial time is given by:[P0 4 JG
MDW

9-3

x

(013CPFSW - 013 CMDW )today
As stated above, the mean phosphate content of today r s deep sea
water is 2.25 l-Im/kg.
We have shown that the amount of dissolved
inorganic -earbon in the sea was about 6% higher during glacial
than during interglacial time and that the amount of water inwhich this carbon was dis sol ved was about 3.5% lower than today.
This yields a ratio of [LC02JG to [EC02Jtoday of about 1.09.
Thus:
2.7
[P0 4 JG
= - - - - x 1.09 x 2.25
MDW
2.1
or

THE COMBINED EVIDENCE FROM DEEP SEA CORES
We can now put all these resul ts together and calculate theC02 partial pressure for surface water during glacial times.
Asshown in table 9-5, we start with the composition of average
present-day deep water (line 1).
We then show (line 2) that if
this water were to reach the surface in a "Strangelove" ocean(i.e., one with no life), then the C02 pressure would be close to
three times the present atmospheric value.
In the real ocean

472

-I
2
Q..------..----,.....-----.,:r~ 2.7 0/00

40

(GLAC IAL).......-..

MEAN DEEP

"-

WATER

~

rt')

0
......., 80w~2. I %0(9

(INTERGLACIAL)

<{

120

PHOSPHORUS
FREE SURFACE
WATER
160

~12. Scrematic
diagram of the glacial to interglacial
1
cycle in the 0 C values for the ECO 2 in mean deep water and for
the EC02 in phosphate-free surface water.

Figure473

however, the activity of plants would deplete the phosphate (andnitrate) and utilize some of the carbon to form soft tissue. The
chemistry of the water after this had happened is given in line 3of the table. The alkalinity change between lines 2 and 3 is theresult of nitrate utilization.
We have one step to go.
Some ofthe plants which live in surface water form CaC03 cages; some of
the animals which eat these plants form CaC03 shells.
The impact
of CaC03 formation is shown in line 4.
He re we mus t digres s a bi t to show how the ratio of CaC 3
formation to plant tissue formation (i.e., the ratio of CaC03 rain
to P0 4 rain)in today's ocean is obtained. We adopt a value of one
mole of carbon lost to CaC03 for each four moles of carbon lost toorganic tissue. This value is obtained from the comparison of thechemical composition of warm surface waters from various regionsof the ocean with that for mean deep water.
If chemical differences between surface water and mean deep water (after normalization
to the same salinity) are attributed to the production of CaCa 3
and organic tissue, then the magnitudes of the removal of eachphase can be calculated. The amount of CaC03 produced is obtained
taking half of the nitrate corrected alkalinity change, i.e.,°

LiCCaC03

9-4

=

2

As we have mentioned before, the removal of N03 ion to form
organic tissue raises the alkalinity of surface water, partly
compensating for decrease in alkalinity resulting from CaC03
formation.
The amount of carbon fixed into organic tissue can be estimated in two ways.
It can be obtained by subtracting the amountof CaC0 3 produced from the difference between the rC02 concentrations for-deep and surface water:

9-5
It can also be estimated by mui tiplying the change in phosphateconcentration by the C/P ratio in organic tissue (i.e., by 105).
An example of such a calculation is shown in table 9-7. The ratio
LiCorg/LiCCaCO
is then calculated by averaging the two LiC org
estimates. the results of a number of such estimates are given intable 9-8. They range from 3.5 to 7.2 and ave rage 4.5. This result must be corrected for the fraction of each phase which accumulates in the sediment. As we have seen only about 1% of the organic tissue survives destruction hence no correction need be
made.
However, in the case of CaC03 about 15% of that produced
accumulates in the sediments.
Hence the LiCCaC03 estimate mustbe multiplied by 1.15.
This brings the ratio LiCorg/LiCCaC03
close to 4. We adopt this value.
While the exact choice of the ratio of LiCorg/LiCCaC03 is
not very important, the extent to which it changes from glacial tointerglacial times is.
Such changes would generate a differencein the C02 pressure for surface water between glacial and interglacial time!
The control on this ratio is ecological.
It
0

474

Table 9-7. Example of the method used to calculate the ratio oforganic carbon to CaC03 carbon produced in surface waters.

N03
l.1m/kg

2:C0 2
l.1m/kg

Alk
l.1eq/kg

Average
Deep Water

2.22

32.1

2245

2354

GEOSECS Station
214 Surf H2 0 *

0.04

0.1

1926

2266

Difference

2 . 1 ts

32.0

319

88

88+32
DoCCaC03 ==DoC O rg

= 60

319-60 = 259

DoCCaC 03=(Method #1)

~m/kg

-

DoC org -= 105x2.18
DoC org

~m/kg

2

229

~m/kg

(Method #2)

244

=

= 4.1
60

* Normalized to salinity of average deep water.

475

Table 9-8. Estimates of the organic to CaC03 carbon fixation rate
for surface waters from throughout the temperate and equatorial
oceaneBased on results obtained as part of the GEOSECS program(424, 425, 426, 406).
T
°C

Stne
No.

bCCaC03~m/kg

Method #1*
tlC orgl.Jm/kg

Method #2*
~Corg
~m/kg

~Corg/ bCCaC 0 3

ATLANTIC

31
37
42
54
61
105
115
120

27
28
27
26
17
25
19
18

34
38
59
61
56
56
55
65

260
298
263
277
229
243
234
240

227
231
224
222
217
210
230
226

7.2
7.0
4.1
4.1
4.0
4.0
4.2
3.6
4.8"

INDIAN

43e
442
446
417
420
427

-

27
30
28
27
27
26

66
65
38
59
54
56

239
254
242
240
227
245

222
221
226
183
209
217

3.5
3.7
6.2
3.6
4.0
4.1
4.2

229
190
228
174
208
225

4.0
4.4
3.5
4.1
5.1
5.1
4.4

PACIFIC

214
201
227
244
260
269

28
18
27
26
29
27

60
37
66
48
47
49

259
133
232
220
291
275

*Note that method #1 gives on the average a 40 ~m/kg higher
answer than method #2.
One reason for this difference is thehigher preformed ~C02 content of cold than warm surface water which
contributes to the difference between the ~C02 content of phosphate
free surface water and cold deep water.
Were the resul ts frommethod #2 alone used then the mean would drop to about 4.0.

476

depends on what fraction of the plants form GaG03 cages and on
what fraction of the animals which eat these plants form GaG03
shells.
As we have at present no way to assess the change inocean ecology from glacial to interglacial time, we assume thatthe ratio of 4 applies to both periods.
Now back to the calculations in table 9-5.
The composi tionof glacial mean deep sea water is obtained as follows.
The P04
content of 3.2 ~m/kg obtained above is used. The N03 to P0 4 ratio
in the sea is assumed to have remained at 15.
As the process
which removed PO q from the ocean at the end of glacial time would
likely have also removed N0 3 it is unlikely that the N03/P04 would
have changed.
If it did change then we would have to calIon theecologic mechanism described in chapter 6 to produce the adjustment. Thus we take the glacial N03 concentration of deep water tobe 15 x 3.2 or 48 ~m/kg.
The mass of EG0 2 in the ocean is taken to be 6% higher during
glacial than during interglacial time and the salini ty of theocean to be 3.5% higher.
Thus the EGO 2 concent ra tion in deep
water is taken to be 1.06 x 1.035 the interglacial value.
There
are two ways in which the alkalini ty might be computed.
One way
would be to make the alkalinity change exactly equal to the EG02
change.
While this is approximately the result expected if thecompensation were complete, it is not the best estimate that canbe made. A better approach is to use the difference in the global
lysocline level between full glacial and full interglacial to fix
the difference between the C03= ion concentration in average
glacial and interglacial deep water.
If the EC02 and C03=
concentrations for glacial deep water are known then the alkalinity is determined.
Studies of the extent of dissolution in cores taken in thetransition zone where calcite is partially dissolved reveal thatwaters in the equatorial Pacific and Indian Oceans were slightly
less corrosive during glacial than during interglacial time (496,501).
Tnose in the equatorial Atlantic appear to have beenslightly more corrosive (497, 507).
Al though the exact level ofthe lysocline during glacial time has yet to be established, it
appears that the glacial to interglacial shifts were on the orderof only a few hundred meters.
In today's ocean a 1.0 to 1.5 ~m/kg
change in C03= ion will produce a lOa-meter shift in lysocline
depth (see chapter 2).
Thus the changes in lysocline observed
between glacial and interglacial time likely require no more than
5 ~m/kg changes in CO 3 = ion content.
We es tima te tha t during
glacial time average deep wa ter had a 3 ± 3 ~m/kg higher CO 3=
ion content than during interglacial time.
Based on this assumption the alkalinity in glacial deep water is obtained.
As can beseen from table 9-5, the alkalinity difference between glacial andinterglacial time calculated in this way comes out to be 212~eq/kg, while the EC02 difference is 216 ~m/kg.
Thus the two
approaches give almost the same answer. This completes the reconstruction of the chemical parameters for the average deep water ofglacial time (line 5, table 9-5).
Were this water brought to the surface in a "Strangelove"ocean (line 6) it would have had an even higher Pea
than its
interglacial eqUivalent.
However, after plants compiete their477

Table 9-9- Sensitivity of the C02 partial pressure in surface waterand to various properties of surface water.
These calculations werecarried out using the apparent dissociation constants of carbonic acidin sea water determined by Mehrbach et al., at Oregon State University(400), and the C02 solubility of Weiss at Scripps Institution of
Oceanography (403) following the procedures outlined by Broecker and
Takahashi (463).
In each example one paramet er is changed and theothers are held constant.

Temp.
°C

Sal.

PO If *

0/00

~m/kg

Alk-EC0 2

C03=

~m/kg

~m/kg

4
4
4
4

315
315
315
315

222
222
221
220

332
319
305
292

4
4
4
4

315
318
320
323

222
224
226
228

332
336
342
346

Alk.
ueq/kg~m/kg~CaC03

2276
2276
2276
2276

1961
1961
1961
1961

2276
2296
2315
2335

1961
1978
1995
2012

EC02

~Org

p~O 2

10- atm

Change Temperature

23
22
21
20

34.7
34.7
34.7
34.7

2.2
2.2
2.2
2.2

Change Salinity

23
23
23
23

34.7
35.0
35.3
35.6

2.2
2.2
2.2
2.2

Change tC02 (hold alk. constant, i. e. , no CaC03 compensation occurs)

23
23
23
23

34.7
34.7
34.7
34.7

2.2
2.2
2.2
2.2

2276
2276
2276
2276

1961
2001
2041
2081

4
4
4
4

315
275
235
195

222
196
171
146

332
404
498
622

Change ECO 2 (hold alk- ECO 2 constant, i. e. , CaC03 compensation occurs)

23
23
23
23

34.7
34_.7
34.7
34.7

2.2
2.2
2.2
2.2

Change 1\0 rg/ I\CaCO 3

23
23
23
23
23

34.7
34.7
34.7
34.7
34.7

2.2
2.2
2.2
2.2
2.2

2276
2316
2356
2396

1961
2001
2041
2081

4
4
4
4

315
315
315
315

222
224
225
227

332
344
357
370

2216
2246
2276
2306
2336

1931
1946
1961
1976
1991

2.6
3.2
4.0
5.4
8.2

285
300
315
330
345

200
211
222
233
244

364
347
332
318
305

1856
1909
1961
2013
2066

4
4
4
4
4

378
346
315
284
252

262
242
222
203
183

237
281
332
393
471

Change PO If . in upwelled water

23
23
23
23
23

34.7
34.7
34.7
34.7
34.7

3.0
2.6
2.2
1.8
1.4

2234
2255
2276
2297
2318

*In upwelled water.

478

task of utilizing the PO lf (and N0 3 ) the peo 2 drops well below
the value for interglaci~l water.
The peo
for the glacialsurface ocean is 80 x 10- atmospheres below tbat for the interglacial surface water.
The ocean sediment record does indeed
generate a result which is consistent with that obtained directlyfrom measurements on ice!
The sensi tivi ty of the pe0.z difference to uncertainties inthe specifics of the reconstructlon is given in table 9-9.
While
it is difficult to assign an overall uncertainty to the glacial
pe02 reconstructed in this way, it is probably no less than ±25
x 1u- 6 atmospheres.

CAUSE OF THE OCEANIC PHOSPHATE CHANGE
What caused the P0 4 content of the sea to drop at the closeof glacial time?
The mean replacement time for PO 4 in the sea( -100, 000 years) is too long to permit us to calIon the steadystate control mechanisms outlined in chapter 6.
Rather, we must
once again turn to the generation and destruction of a small
phosphorus bearing reservoir.
Of the small reservoirs discussed
above, the terrestial ones (soil and trees) are of no use.
Phosphorus cannot be transferred through the atmosphere from sea tocontinent. Hence while the growth of trees and the development oforganic-rich soils during early post-glacial time could bringabout a change in the }";e02 content of the ocean, the enlargementof these reservoirs would not lead to 'any change in the phosphorus
content of the ocean!
On the other hand, the loss of carbon to
shelf sediments would be accompanied by a loss of phosphate (andnitrate).
It is interesting to note in this regard that the ratioof e atoms lost to organic matter to phosphorus atoms removed fromthe sea at the ct~se of glacial time needed in order to produce nochange in the 0 e for planktonic foraminifera is close to 105.
Thus if ttie organic matter deposited on the shelves contained carbon and phosphorus in roughly their "Redfield" ratio, then shelfremoval could explain what we observe!
So we come back to a loose end we dropped earlier in thischapter.
What were the relative contributions of terrestrial
reservoirs (trees and soil) and marine reservoirs (shelf sediment)
to the removal of carbon from the ocean-atmosphere reservoir at
the close of glacial time?
Shackleton makes a good case that atleast part of the loss must have been to the terrestrial reservoirs.
Let us attempt a corresponding estimate of how much material might have been deposited on the shelves during early postglacial time.
We know from engineering borings that the world's
estuaries received several tens of meters of sediment during earlypost-glacial time.
This sediment rebuilt the sedimentary ramp
from river to the shelf edge destroyed by erosion during the low
sea stand of glacial time.
While similar deposition is likely to have occurred on many
parts of the shelves eroded during glacial time, it is difficult
to assess the magnitude of these deposits.
We can, however,estimate roughly how large they would have to be to explain thebenthic o13 e shift.
If we assume that the average shelf sedimentcont ains 1% by weigh t organi c carbon, then the mas s of these479

sediments would have to be 100 x 9 X 10 16 moles x 12 g e/mole or 1
x 10 20 g. The area of the world's shelves is about 3 x 10 17 cm 2 •Thus if this sediment were spread uniformly over the entire shelfarea it would make a deposit with 360 grams of mineral matter/cm 2
or a layer three to four meters thick.
While such a mass is not
so large that it would cause a person familiar wi th shelf geomorphology and sedimentation to flinch, such a person might bemore comfortable with an amount 3 or so times smaller.
Thus, asis the case for trees and soils, we can only show that shelf sedimenta tion could do the job.
Hence the loose end remains.
BothtEJrrestrial and marine reservoirs must have contributed to theo 3e change observed for benthics but it is not possible to state
which played the more important role.
A compromise could be made by assuming both terrestrial andmarine contributions to the organic carbon removal and a lower
than Redfield e/P ratio in the organic material accumulating on
the shel ves.
In the above discussions we have assumed that theorganic residues falling into the deep sea and the organic residues accumulating in shelf sediments have e/P ratios of close to105.
However, these elements might be differentially released
during decomposition.
Indeed measurements of the organic e to P
ratio in shelf sediments range all the way from 50 to 250.
As
these sediments are extremely heterogeneous, it is difficult tocome up with a meaningful average.
AN ALTERNATE SCENARIO

The possibility of substantial deviations between the chemical composition for the particulate debris falling into the deepsea and the standard Redfield ratios opens a whole new eXPt~­
nation for the difference between glacial to interglacial 0 echanges for benthic and planktonic shells.
The 13 e record could
be equally well explained by a 30% glacial to interglacial decrease in the e/P ratio for the organic debris falling from surface to deep water.
In chapter 3 we demonstrated that there is adifference in the e/P ratio of material oxidized from place toplace in today' s deep ocean (see table 3-5).
Thus a 30% change
from glacial to interglacial time is not beyond the realm of
possibili ty.
A 30% reduction in this ratio would cause the observed 0.6 0 100 reduction in o13 e difference between surface anddeep wa ter ~eo z.
If this change were coupled wi th storage oforganic matter in trees and soils then both the o13 e history forplanktonics and benthic forams could be explained.
In fact,
everything else about which we have spoken (the calcite compensation, the peoz change, •.. ) would be the same as for the shelfstorage hypothesis.
So the plot thickens!
We now have not only
the relative importance of terrestrial and marine organic carbon
storage to worry about, but also the relative importance ofphosphorus removal to shelf sediments and of a change in the e/P
ratio in organic debris falling to the deep sea.
Is there any way to get a handle on the latter? Perhaps!
To
understand how, we must consider the sequence of events at theclose of glacial time.
To simplify things we will contrast thetwo extreme scenarios: i.e., organic carbon and phosphorus storage480

in marine sediments on one hand, and terrestrial storage oforganic carbon coupled wi th a change in the C/P ratio in fallingorganic debris on the other.
These endmember scenarios are shownin figur 9-13.
Also shown are the corresponding time historiesof the 0Y3 C changes corresponding to each.
While the benthic rI~ords should be basically the same (aunidirectional rise in 0 C during early post-glacial time) the
planktonic records might be different in one respect. In the caseof the residue composition scenario it is £ossible that a "spike"might have occurred in the planktonic 0 1 C record.
This spikewould be produced by the lack of synchronei t{ between the change
in ocean ecology (which would decrease the 0 3C for planktoni1~)
and the growth of forests and soils (which would increase the 0 C
for planktonics).
If, for example, the ecology change occurrred
before the fores t and soil development, then the spike would be
toward low 013 C values (see figure 9-13).
The duration andmagni tude of the spike would depend on extent of separation intime of the ocean ecology and terrestrial storage events.
In nocase, however, could the magnitude exceed 0.7 0 /00.
No suchspike would be possible for the shelf storage scenario because theremoval of carbon and phosphorus to the shel ves would take place
simultaneously.
Is there any evidence for a spike in the planktonic 0 1 3C
values in early post-glacial time?
As was the case for the calci te preservation event a systematic search has yet to be made.
As shown in figure 9-14, the results on cores run to date show nosign of such a spike. While the record for core V12-122 from the
Caribbean Sea shows an anomaly during Holocene time, this anomaly
is seen for sacculifer and not for ruber. Also this record is so
noisy that what appears to be an anomaly may well be just anotherrandom fluctuation.
The more detailed records for cores EROC 92
and l2e from the western equatorial Pacific do not show suchanomaly. In examining these records it must, however, be kept inmind that bioturbation would take a heavy toll on an event only a
few thousand years in duration. Until more detailed studies havebeen made on cores with higher deposition rates this question mustremain open.
WRAP UP OF THE CD2 RECORD

Summarizing, then, it has been shown that at the close of
glacial time about 3% of the carbon in the ocean-atmosphere system
was removed. as organic residues. The removal of this carbon triggered a CaC03 preservation event in the deep sea. Further, either
the phosphorus to carbon ratio for mean ocean water decreased
through phosphorus removal to shelf sediments, or the phosphorus
to car bon rat i 0 in the 0 r gani c deb r i s r a i n i ng t o t he dee p sea
decreased by about 30%.
This reconstruction yields a glacial tointerglacial increase in the CO 2 partial pressure for surface water (and hence also the atmosphere) close to that observed in the
air bubbles trapped in polar ice.
It was not possible, on thebasis of the information available when this book was written, todistinguish between the roles of the terrestrial reservoirs, of
shelf reservoirs, and of ocean ecology in producing these changes.

481

a l3 c PFSW(%o)
2al3 c PFSW (%0)
GLACIAL
CONDITIONS(25,000 TO 13,000)

3

•

t

I
81

I

T
I
M

E

~

p

,~

~

BI

1

I

1

~

I

FOREST GROW TH
SOIL DEVELOPMENT(11.000 TO 8000)(12,000 TO 7000)

I

I
I

L.+

I

/

I
l

./

T
I
M
E

~~

CALCITECOMPENSATION+IN DEEP SEA(-10,000 TO -4000)

~

I
I
I

0

~p

CHANGE IN OCEANECOLOGY AND IN C/PRATIO IN ORGAN ICRESI DUES (-IZ,OOO

SEA LEVEL RISEDEPOSITION OF
ORGANICS ON
SHE LVES

3I

GLACIAL PERIOD...- COMES TO AN END(- 12. 000)

I

-I

2

I

INTERGLACIAL
CONDITIONS(-4000 TO PRESENT)

I

I

-

0

I

al3 c MOW (%0)

8 13C MOW (O/oo)

Ft~ure 9-13.
C;
C record

Comparison of two scenarios which could explain thefor planktonic (P) and benthic (B) forams and thepca 2 record obtained from ice cores.
The records differ only inthat the right hand scenario might have produced a brief spike in
the C;13 C value for planktonic forams in early post-glacial time.

482

,
.,
-r80-t. .

al3 c (%0)

a'60 (%0)

o -2

fR~e

I

161°E
m3100m';';

-

~.

...

a~o·s
128

10 f-

e

-

1

-I

10

~

20 -

-

000

o
~00

40 -

G. SACCULlFER• OXYGEN.

50
0

E

10

Sl80 (%0)
-2

-I

0

fROe

o 92

8 2·SI57°E
<:)

~

u

I 20
t£t.

1600m

9

1" L

0
00-

W

o 30

TERM I

Cb'b

G. SACCUL IFE R

OX'rGEN

••

•

~

-

I.'

&
£t.
0
...
W 30 -TERMI g -

J:
t-

2

I

-t-. .:-• -- -..-. -

_

-.
,

°

.,'f

\

_.

-.--.-

•

~

•

•

~
I

2

f

•
•

-

.

-

t
J

G. SACCULlfER
CARB.ON

• 13,

I

~-

-

•• •

.

-

G. SACCULlfER
CA BON.•-~_., ~-•300
TERMII-.

G. SAC CU'

Po OBUaullOCULATACARIBONOXYGEN

320

40

18

•
-•
•
,,[~l
• •

13

G. SACCULIFEk
CARBON••••

G. RUBE R
CARBON

Figure 9-14. The records of <5 0 and <5 C for planktonic forams
in two equatorial Pacific cores and in one Caribbean core.
There
is no evidence of a <5 1 3C spike in early post-glacial time.
The
results on EROC 92 and 128 were obtained by Berger and his coworke rs at·. Scripps Ins ti tu tion of Oceanography (329); those oncore V12-122 we re obtained by Van Dank at the Lamont-Doherty Geological Observatory (327).

483

However, more detailed studies of the records in ice and in marine
sediments should yield addi tional evidence needed to narrow down
the range of possibility.

THE OXYGEN RECDRD
If during glacial time the phosphorus to carbon ratio ineither mean ocean water or in particulates falling to the deep sea
were larger than today's then the dissolved oxygen content of the
deep sea water would have to have been lower.
The reason is asfollows.
As the temperature of deep water in the sea during glacial times was only 1.5 ± 1.5°C colder than today's, the solubility of oxygen in the descending glacial deep water would have beennearly the same as in today's. Further, the oxygen content of theatmosphe~e could
not have been appreciably different than in
today' s.
The amount of dissol ved oxygen brought to the deepsea wi th each uni t of new deep wa te r would als 0 have been thesame.
On the other hand, the demand for oxygen in the deep sea
would have been greater.
If there were more phosphorus dissolved
in glacial deep water, then for each unit of deep water upwelled
more phosphorus would have reached the surface and more organic
material would have rained to the deep sea.
If, on the other
hand, the particles raining to the deep sea during glacial time
carried more carbon per uni t of phosphorus than those falling
today, the oxygen demand would also be hi~her.
Thus, regardlessof which explanation is adopted for the 0 1 C results, a lower dissolved oxygen content for glacial deep water than for interglacial
deep water is required.
While the amount of the decrease in oxygen in glacial deep
waters depends on the details of how the ocean circulated during
glacial time, a rough es tima te can be made from the P04 content
difference between glacial and interglacial deep water required by
the shelf storage hypothesis.
As listed in table 9-5 thephosphate content of mean deep water dropped by close to 1 lJffi/kg
at the close of the glacial time.
Using the ordinary Redfieldratios, this leads to a difference in the oxygen content of deep
water of 135 \.lm/kg.
As the mean oxygen content of deep water is
now only about 150 l1m/kg, such a reduction would by no means be
minor.
Is there any way to read paleooxygen contents from the record
in deep sea sediments?
In chapter 6 we discussed one possible
method. It. appears that the shells of the same benthic foraminratios
ifera that record ice volume through their 18 0 /16 0*As the trean replacement time of oxygen in the atmosphere iscurrently on the order of 10 million years, no significant changeon the time 16scale of a few thousand years could have 0r~urred.
The 12 x 10
moles of 02 made available when the 9 x 10
moles
of carbon were stored as organic residues duri~% early post
glacial time is trivial compared to the 15,000 x 10
moles of 02
already in the atmosphere.
Finally, experimental evidence exists
for this constancy. The 02/N2 ratio in the air bubbles from iceof glacial age is identical within the measurement error to thatin today's atmosphere.

484

and the mass of stored organic material through their 13 C/ 12 C
ratios may also record dissolved oxygen concentration by their
species abundance.
Studies of the benthic communities in deep
Atlantic sediments by Streeter at the Lamont-Doherty GeologicalObservatory (619) show that those species whose abundance isgreater where low-oxygen bottom waters bathe today' s sediment,
dominated in abundance over a wider area of sea floor during
glacial time than during Holocene time.
Correspondingly, thosespecies found in abundance in sediments bathed today in highoxygen waters had more restricted areas of dominance during glacial time.
As similar studies have yet to be carried out in
Pacific and Indian Ocean sediments even a qualitative inventory ofoxygen contents of mean deep water cannot yet be attempted.
Allthat can be said at this point is that the Atlantic resul ts are
consistent with the other chemical changes which appear to havetaken place between glacial and post-glacial time.

GLACIAL TO INTERGLACIAL CHANGES IN OCEAN MIXING RATE
So far we have devoted all our efforts in this chapter toglacial to interglacial changes relating to the change in atmospheric C02 pressure suggested by measurements on ice cores.
Other changes may have occurred which are not tied to the transfer
of C and P to and from small reservoirs or to changes in the elPratio in raining particulate debris.
Is there, for example, anyevidence that the rate of oceanic mixing was different during
glacial than during interglacial time?
One way to get at thisproblem is to look for a change in the rate of accumulation of
CaCO 3 above the lysocline.
Other things being equal, a higher
mixing rate should produce a higher calci te rain rate.
As thepreservation event discussed above would be felt only below thelysocline, we need not be concerned with its effects.
The pes ul ts from cores on which appropriate 14C dating andcalcite percentages have been measured are summarized in table
9-10.
In all three cores CaCa 3 accumulated fas ter during lateglacial time than during post-glacial time.
However, because thecores analyzed
represent only a few spots in the equatorial andsUbtropical ocean it would be premature to claim that they are
indicative of the entire ocean.
Although the calcite compensation spike would not confuse
attempts to estimate the mean ocean mixing rate from super lysocline CaCO 3 accumulation rates, one other aspect of the scenariopresented 'above might confuse the situation.
If there were 30%
more phosphorus present in the ocean during glacial time then thesame rate of mixing would produce 30% more plant growth and presumably 30% more CaC03.
Hence, if the phosphate content explanation for the pcachange is adopted then a somewhat higher
glacial rain rate o~ CaC03 would be expected.
Perhaps this 1s the
explanation for the results on two of the cores in table 9-10.

GLACIAL TO INTERGLACIAL LYSOCLINE CHANGESThe lysocline of the ocean may ha ve undergone glacial tointerglacial changes independent of the preservation event we have485

Table 9-10. Calci te accumulation rates for late glacial and for
Holocene time as determined by radiocarbon dating of cores taken
above the lysocline.

Glacialg
cm 2 10 3yr

Holoceneg
cm 2 10 3 yr

Ratio

Ref.

lan

61°E

3.4

2.4

1.7

1.4

( 197 )

10 0 S

24°W

3.3

3.7

1.6

2.3

(195)

2°S

157°E

1.6

2.3

1.8

1.3

( 199 )

Core
No.

Lat.

V19-188

7°N

A180-74
ERDC92

Long.

Depth

486

discussed.
For example, if the rate of ocean mixing was higher
during glacial time and if this change induced a higher CaC03 rain
rate then it should have produced a change in the steady state
depth of the lysocline.
As stated in chapter 6, the level of the
global lysocline must vary wi th the rate of oceanic mixing andwith the rate of carbon input from rivers.
The adjustment time
would be several thousands of years.
Thus during the same periodthat the CaC03 preservation spike was in progress, the oceanic
lysocline may also have been simultaneously adjusting to the newinterglacial CaC03 budget.
If our understanding of the history of the lysocline is toimprove we must develop a more quantitative scheme for assessing
the extent the calcite in any given sediment has been dissolved.If
this could be done both accumulation rates and rain rates could beobtained.
Although the methods described in chapter 6 provide
indices of calcite dissolution, these indices remain uncalibrated.
Let us consider a method by which this might be accomplished.
The approach has to do with the radiocarbon age of thecalcite in the core-top bioturbated layer.
We have already mentioned that this age should be approximately equal to the thickness of the mixed layer divided by the sedimentation rate.
We didnot mention that this approximation holds only in the absence of
calcite solution.
Let us now do the job rigorously including thepossibility that solution occcurs.
Figure 9-15 shows the fluxes
invol ved.
Radiocarbon is added to the bioturbated zone wi th theraining calci te.
It is lost from the bioturbated zone by radiodecay, by burial, and by calcite solution.
For steady state theequation relating these fluxes is as follows:C

9-6

The symbols in this equation are defined in figure 9-15.
Thi s
equation can be solved for the ratio of the 14 C contents of mixed
layer and ··surface ocean carbon:
1

1 + fphAIR
This ratio is related to thefollowing equation:

14 C

9-7

age for mixed layer carbon by the

9-8

487

14 C
--t-....-.(I-fp)fRRC}M

SOLUTION

R = TOTAL DEPOS ITION RATE

A: TOTAL ACCUMULATION RATEfR: FRACTION CALCITE IN RAINING MATERIAL

f p = FRACTION CALCITE PRESERVED
f S = FRACTION CALCITE IN SEDIMENTh = DEPTH OF BIOTURBATION

A= DECAY CONSTANT OF 14C
14/}O= RATIO IN SURFACE OCEAN rC02
14

--.f)
·C

= RATIO IN MIXED LAYER CALCITE
M

Figure 9-15. Flux balance for radiocarbon in the bioturbated layerof a marine sediment undergoing partial calcite dissolution.

488

Eliminating the activi ty ratio by simul taneous solution ofthese two equations and solving for t we get:t =~ In [1 + f:hA ]9-9

Substituting the mean life of radiocarbon TM, for l/A, we get:

9-10
At this point it should be noted that if f p is unity then:

9-11
For h/A «TM this equation can be approximated as follows:
h

9-12

t

A

This is the approximation we used earlier for the age of CaC03 inthe bioturbated layer in cores not subject to CaC03 solution.
In situations where calcite solution occurs, R, the rain rateof sediment, need not equal A, the accumulation rate of sediment.
Further the rain rate is not known (the accumulation rate can bemeasured by 14 C dating beneath the bioturbated layer).
Hence wemust establish the relationship between R and A in terms of thefractions of fp and fS. This relationship is as follows:
A

9-13

(1

where (1fR) R is the accumulation rate of noncalcite andfpfRR is the accumulation rate of calcite.
Unfortunately fR
is also an unknown.
It can, however, be related to fS and fp
by the following equation:

9-14
where the numerator is the fraction calci te in the sediment andthe denominator is the sum of the fraction calcite and thefraction noncalcite in the sediment. Solving for fR we obtain:

fp + fS - fpfS

9-15

Now by combining equations 9-10, 9-13 and 9-15 an equation relat-489

Table 9-11. Estimates of the extent of solution of calcite based on

the radiocarbon method for two cores from above the lysocline.

Core #1 ERDC 92 (2°S, 157°E, 1.6 km) EQUATORIAL PACIFIC OCEAN (199)h = 8 cm

A

fp
fp
fp

1.0
0.9
0.8

4890 yr
4200 yr
3520 yrtttfS = 0.86

tobs = 4300 ± 300
i.e., fp = 0.92 ± 0.05

Core #2 V19-188 (7°N, 61°E, 3.4 km) EQUATORIAL INDIAN OCEAN (197)
2.0 cm/10 3 yrs

A

t

=

8.27 x 10

3

1.0
0.9
f p =. 0.8
fp
fp

~n

h = 9 em

9f p

[1 + 2.0 x 8.27 (1 - f
t
t
t

=

0.84(1 p +f s = 0.84
f p)

)]

0.84(1 - f p )tobs = 3200 ± 200
i. e. , f p = 0.93 ± 0.04

3593
3060
2540

490

ing to t and fp can be wri t ten in terms of measurable parameters(h, fS and A). The result is:
t

=

TM In

[

1 + fphTMA(1-

fS (1 - fp)

) ]

9-16

fp + f S (1 - fp)

While this equation is not conveniently solved for fp a graphical solution can be made (i.e., calculate t for various values offp and then plot one variable against the other).
To date the information needed to apply this method exists
for only two cores. Both are from above the lysocline. As can beseen in table 9-11, both yield preservation factors for calcite(i.e., values of fp) close to unity.
The 14 C results support
the assumption made throughout the book that minimal solution
occurs above the lysocline.
The next step is to examine coresfrom beneath the lysocline.
The idea would be to use the fp
values obtained from 14 C data to calibrate dissolution indices(ratio of foraminifera fragments to whole tests, the ratio of
whole planktonic to whole benthic tests, •.• ).
Once calibrated
these indices could be used to quantify the change in the extent
of calcite solution between glacial and interglacial time.
There is, however, a potential pitfall with this method.
Itlies in the implicit assumption that all the grains of calcite
lose the same fraction of their calcite each year.
We know that
this cannot be the case.
The fact that the shells of certain
species of foraminifera lose their identity through solution-induced breakup into fragments faster than others suggests that thesolution process is not homogeneous.
The way to check the impactof this assumption would be to perform radiocarbon dating on various size fractions (i.e., fragments versus whole forams) and evenon hand-picked specimens of a single species. While the latter iscurrently -impossible because of the sample size required for 14 C
dating (i.e., 3 to 10 g of CaC03), the recent invention of thehigh energy mass spectrometry technique* will reduce the sample
size by a factor of 1000.

CRANGES IN THE DISTRIBUTION OF NUTRIENTS IN THE DEEP SEA
In our discussions of glacial to interglacial changes inocean chemistry, we have thus far spoken only of the mean composition of deep water. We must also consider the gradie~n theseproperties 'from the North Atlantic through the Antarctic into theNorth Pacific.
Studies of the sedimentary record in the northern
Atlantic suggest that the source of North Atlantic Deep Water (atleast as it now exists) was cut off during much of glacial time.
If so, the pattern of deep sea ventilation during glacial timemust have been qui te different from today' s .
So also would the
distribution of nutrients.
How do we know that the present sources of northern component
water was cut off during glacial time?
One such line of evidence*This technique measures the 14C atoms themselves rather than theS- particles they emit during radioactive decay.

491

comes from paleontologic data.
In today's ocean the position of
the Gulf Stream is clearly seen in the sediments. Because of thedifference in the organisms living within and to the north of thestream, the material raining to the sea floor changes as theboundary is crossed.
The same contrast can be seen in glacial
sediments but the course marked out is quite different (see figure9-16) .
During glacial time the stream left the margin of theAmericas at roughly the same place as it does today but rather
than moving to the northeast as it does now, during glacial time
it went directly to the east (i.e., toward Spain rather than toward Iceland).
The Gulf St ream's Pacific Ocean counterpart, theKuroshio, has at present an eastward rather than a northeastward
trend.
To the north of the Kuroshio lies a zone of lower surface
water salini ty.
Upwelling occurs in. this zone.
It is thoughtthat a similar si tua tion may have exis ted in the Atlantic during
glacial time.
Another line of evidence comes from the oxygen isotope record
for benthic foraminifera in the Norwegian Sea sediments (see figure 9-17).
The deep water in the Norwegian Sea is now about 4°C.
colder than that in the open Atlantic.
This difference is reflected in the 6 l 8 0 values for shells of Holocene benthic foraminifera in th
two basins.
For benthics deposi ted during glacialtime, the 61 8 0 difference changes sign suggesting that the Norwegian Sea waters were slightly warmer than those in the open
Atlantic!
If so then the deep Norwegian Sea must have been ventilated by waters spilling in from the North Atlantic rather than
through the sinking of Norwegian Sea surface waters.
This evidence shows dramatically that the North Atlantic deep water source
as we know it today was not functioning during glacial time.
If indeed NADW was not flowing during glacial time , it ispossible that all three major ocean basins were ventilated by deep
wa ter formed in the Antarctic (as the Pacific and Indian Oceans
are today).
If so, then the gradient in deep water properties
from the North Atlantic around to the North Pacific is not likely
to have been as pronounced during glacial time as it is today.
There are three ways to assess this contrast from thesedimentary record.
First, the 6 13 C: difference between the deep
waters in the North Atlantic and those in the equatorial Paci~~cnow averages about 1 0 /00 (see figure 6-12).
Benthic foram 0 C
data might then be expected to show a smaller Atlantic to Pacificcontrast in 0 1 3C during glacial than interglacial time.
Second,Atlantic deep waters now have considerably higher oxygen contents
than Pacific waters. Oxygen-sensitive benthic species should showless contrast in their distributions between the Pacific andAtlantic during glacial than during interglacial time. Third, thelysocline now stands deeper in the Atlantic than in the Pacific.
The levels should have been more nearly equal during glacial time.
Wha t does the re cord show?
No paleooxygen es tima tes are as
yet available for the Pacific Ocean.
As c~q be seen from table
9-1, the glacial to interglacial change in 6 C for benthic forams
living in the Atlantic averages slightly more than the change forbenthic forams from the Pacific.
While this is consistent with aglacial reduction in the interocean difference during glacialtime, more cores will have to be run before the difference can be

492

80 0

60 0

./'""'
I

9-16. Maps showing the boundaries of various faunal assemblages found in sediment of Holocene age (above) and for the peakof the last glacial time (below).
A striking difference between
these maps is seen in the North Atlantic. The assemblage boundary
which follows the present course of the Gulf Stream went straightacross to Spain during glacial time.
The white areas designateice (glaciers on the continents and winter sea ice in the oceans).
These maps were prepared by members of the CLIMAP group (609).
Figure493

al8 0 (%0)

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Figure 9-17. Comparison of 18 0 /16 0 records for benthic foraminifera in cores taken at 25°N, 44°N and 72°N in the Atlantic Ocean.
As can be seen the nortpernmost of these cores fails to show any
significant change in 0 8 0 value between glacial and interglacialtime.
The -2 0 /00 decrease in 0 18 0 which occurred throughout theworld ocean as the result of the melting of continental ice sheets
appears to have been countered in the Norwegian Sea by a majorcooling (-8~C) of its deep waters at the close of glacial time.
This observation by Duplessey and his coworkers in France (618)
led them to conclude that the Norwegian Sea was ventilated fromthe Atlantic during glacial time much as the Bering Sea is currently ventilated today by waters from the deep North Pacific(618) •

494

9-18. Maps showing the location of opal-rich sediment in
the Antarctic
Ocean during late Holocene time (left) and during
full glacial time (right).
These maps were prepared by Jim Haysand his coworkers at the Lamont-Doherty Geological Observatory(523).

Figure495

demonstrated to be significant.
Finally, we have already pointed
out that the lysocline appears to have been deeper in the Pacificand Indian Oceans during glacial time and shallower in the
Atlantic during glacial time.
However, the magnitude of thisreduction in contrast between the levels in the two oceans has yet
to be established.
Another approach to the problem as to how the pattern of deep
circulation changed from glacial to interglacial time is to compare the patterns of opal accumulation during these two climatic
regimes.
Hays and his coworkers at the Lamont-Doherty GeologicalObservatory have done this for the Antarctic Ocean.
As can beseen in figure 9-18 the belt of opal rich sediment was locatedfurther off the Antarctic continent during glacial time thanduring Holocene time.
A likely explanation is that the floating
ice which now surrounds Antarctica was more extensive during
glacial time pushing the upwelling zone further to sea.

SUMMARY
The first chink in the ocean's armor of secrecy regarding its
chemical history has been found.
Measurements of the C02 content
of air trapped in ice and of the 13 C/12 C ratio in the CaC03 of
foraminifera shells seem to be telling us that the partial pressure of carbon dioxide in the atmosphere was lower during glacial
time and that this change is related to changes in the ocean's
nutrient constituent cycles.
Specifically the ratio of P0 4 to
LCO 2 either in mean ocean water or in the particula te rna t ter
raining from the surface to the deep sea must have been higher
during glacial time than today.
We also have evidence that the circulation pattern of waters
in the deep sea was not the same during glacial time as it istoday.
The flow of deep water originating in the North Atlantic,
which is responsible for large differences in the deep water nutrient element concentrations between the oceans, appears to havebeen cut off during glacial time.
Although confirmation that theflow pattern was different based on paleo 13C and paleooxygen recons truction has yet to be obtained, the potential is certainly
there.
The task of carrying these studies much further back in time
will prove very difficul t.
The greater the time span, the moreconstituents which might vary.
The greater the time span, themore difficult the sedimentary record is to read.
But the challenge is there and marine scientists will surely make concerted
attempts to find more chinks in the ocean's armor of secrecy.

PROBLEMS
9-1

Assuming that 1 fhe hypothesis gi ven in the text is
valid,
construct a 0 C versus P0 4 concentration plot for the water
masses of the glacial ocean.
Assume that the pattern ofcirculation was the same as today's.

9-2

If the C03= ion content in the deep sea during glacial time
was 5 ~m/kg higher than it is now in all the water masses of496

the ocean, how much deeper would the lysocline in the westernequatorial Atlantic and in the eastern equatorial Pacifichave been?
9-3

If the phosphorus-induced 02 reduction in deep ocean oxygen
content was everywhere about 125 llmlkg during glacial time
what areas of the sea floor below 3 km would have been bathedin anaerobic water at that time? How would this affect the
distribution of benthic forams?

9-4

Fossil benthonic and planktonic forams are recovered from ah~)izon dating 20 million years in a deep sea boring.
The<5
C~ value
for the benthics is 0.0 per mil and, for theplanktonics is 4.0 per mil.
If the photosynthetic fractionation factor and the C/P ratio in marine organisms were thesame 20 million years ago as they are now, by how much did
the C/P ratio in the ancient deep sea differ from the C/Pratio in today's sea?

9-5

Consider a hypothetical earth with no glaciers.
This earth
has 30% continents and 70% oceans.
It is at hydrostatic
equilibrium. As glaciation occurs, 4% of the surface area isburied by ice sheets three kilometers thick.
Were hydrostatic balance to be reestablished by how much would the rocksurface beneath the ice be depressed? By how much would thesea level fall?
Geologists living in the next ice-free period attempt to assess the extent of ice formation by studying submerged shorelines.
By how much would the elevationsof shorelines for continental margins differ from thoseflanking islands rising from the ocean floor?
fssume themean density of ice to be 0.9 and of rock 3.3 g/cm .

9-6

A biologist builds a two reservoir aquarium.
One reservoir
has a volume equal to 10% of the total and is exposed to thelight.
The other has a volume equal to 90% of the total andis kept in darkness.
The Ii t reservoir is also exposed tothe atmosphere.
Wa ter is ci rcula ted back and forth between
the reservoirs.
The dark reservoir has the following
chemical composition:
2030 x 10- 6 moles/kg[ LCD 2J
[HC0 3-]1900 x 10- 6 moles/kg[ C03=J
100 x 10- 6 moles/kg
6[ CO 2]=30 x 10- moles/kg
6[Alk]
2100 x 10- moles/kg[N0 3 ]=
30 x 10- 6 moles/kg[P0 4 ]
2 x 10- 6 moles/kg
The <5 13 C value for the LC02 in this reservoir is 0.0 0 /00.
The surface reservoir has no measurable P0 4 and N01~ thesame a1kalini ty and a 10% lower LC02 content.
Its <5 C is
2.0 0 /00.
What is the equilibrium constant for the reaction497

in the dark tank? What must be the photosynthetic fractionation factor?
Calculate the concentrations of C02, HC03- and C0 3== inthe surface reservoir (i t has the same temperature and P as
the dark one). Remember:== [C0 3=J + [ HC0 3-J + [ C0 2J== 2[ C0 3==J + [ HC0 3-J
The oxygen content of the Ii t reservoir is 350 wn/kg.
Wha t
is the oxygen content of the dark reservoir?
Now the biologist adds enough P0 4 and N0 3 - to raise theconcentration of these species in the dark reservoir to 4
1.Jlll/kg and 60 ~m/kg respecti vely.
He observes more plant
growth and that the P0 4 and N03 content of surface waterremains very low.
What will be the new steady state CO 2,
HC03- and CO~; concentration in the lit reservoir?
Whatwill be the 0 C for the ~C02 in this reservoir?
SUPERPROBLEM #9

Turk Reekian, a geochemist, becomes interested in the chemical changes associated with the great extinction of marine organisms which marked the abrupt close of Cretaceous time. He adopts
the computer aquarium of Stover and Eas tland (see superproblem#7). He first calculates the steady state 13C/12C distribution inthe ~C02 of the four reservoirs.
In so doing he assumes that thefractionation during photosynthesis is 20 0 /00.
He assumes thatthe mean LC02 content of the combined reservoir water is 2.1
moles/m 3 and that of P0 4 is 2.lxlO- 3 moles/m 3 • He takes the mean
013 C of the computer ocean's carbon to be 0.0 0 /00.
He assumes
complete ~ecycling of carbon and of phosphorus (i.e., he allows nothroughput- of ei ther element).
Reekian then abruptly cuts off the computer "life" (i.e., thephosphorus reaching the surface reservoirs is no longer converted
to particuate debris).
He then calculates the time trends for
PO 4, 0 1 3 C and PC0 2 in the surface rese rvoirs .
What does hefind ?
What would be the ratio of the atmospheric CO 2 partialpressure for the new steady state to that prior to the extinctionof life?
Once steady state is achieved Reekian turns on the river input.
Eaclf year they carry amounts of carbon and of phosphorus
equal to one part in 10 5 of that in the computer ocean and an
amount of calcium equal to one part in 10 6 of that in the computer
ocean. He also arranges things so that the concentrations of Mg,
K, Na, S04 and Cl remain unchanged.
The initial ratio of Ca to C
in the ocean is 5.
Develop a scenario for the chemical changes
which would occur over the next million years assuming that theocean remained abiotic. In particular how would the surface waterC02 partial pressure change? Assume that when the ocean reaches astate of la-fold supersaturation with regard to calcite that inorganic precipitation of CaC03 would match the input from rivers.
Reekian realizes that wi thin a few thousands of years after the
Cretaceous catastrophe, life once again flourished in the ocean

498

but is interested in speculating about the chemical conditions ina lifeless ocean.
There are of course some intangibles in such a
calculation.
State what they are and how Reekian might have
handled them.

LAST MINUTE ADDITIONS (SEPTEMBER 1982)
Froelich and his coworkers at Florida State have recently
suggested that the GelSi ratio in opal may prove to be a valuablemeans of accessing changes in the oceanic silicate cycle during
glacial time.
This prediction is based on their finding that the
GelSi ratio in the hot waters emmanating from ridge crests is considerably higher than that for mean river water.
If so, glacial
to interglacial changes in the amount of silicate entering the sea
from rivers or in the residence time of silicate in the sea would
produce changes in the germanium content of sedimentary opal.
Boyle and Hester at M.l.T. have recently shown that the CD to
C.a ratio in benthic foraminifera is proportional to that in sea
water (712). As we have seen in chapter 4, the Cd to P0 4 ratio intoday t s ocean is everywhere the same.
Hence measurements of the
Cd content of benthic foraminifera from various parts of the ocean
for a single time horizon should provide a picture of the P0 4 distribution (and hence the circulation pattern) in the sea at thattime.
Boyle, working with Keigwin at Woods Hole, has shown thatduring glacial time the North Atlantic Ocean had about twice its
present CdlCa ratio, strongly suggesting that production of deep
water in the northern Atlantic was greatly reduced at that time(711) •

499

ehapte~/O
CAN MAN OVERRIDE THE CONTROLS?
THE BUILDUP OF FOSSIL FUEL CO2 IN THE ATMOSPHERE AND OCEANSINTRODUCTION
In the last chapter we saw that the great climatic fluctuations of the last million years had an impact on ocean chemistry.
The cycles of C, P, N, and 02 all appear to have changed in tune
with the advances and retreats of the ice.
Man has the potential
to make changes in climate and ocean chemistry larger than thosewhich took place between glacial and interglacial time.
Thiscapabili ty is a byproduct of our need for fuel.
We are now, andwill be for many decades to come, dependent on the energy derived
from burning of coal, oil, and natural gas.
A major byproduct of
this combustion is CO 2 gas.
The elevated CO 2 partial pressure in
the atmosphere generated in this way will warm the earth's surface.
A decrease in the amount of carbon stored in forests and insoils, as the result of man's need for agricultural land and wood,
has also added C02 to the atmosphere.
This CO 2 will" gradually betaken up by the ocean, and will al ter it's chemical cycles.
Ourgoal in this chapter is to use the information we have compiled in
the earlier chapters of this book to develop a means for predicting the partitioning of this excess C02 between air and sea.
Because we do not know exactly how much coal, oil, and natural gas
will be burned nor at what pace we will conduct this combustion,and because we do not know what new pressures will be placed on
our forests and soils, the model we develop will have to be adaptable to a wide range of CO 2 production scenarios.
C02 PRODUCTION IN THE PASTThe first step in our analysis is to look into how much C02
we have produced to date.
The amounts produced by fossil fuelburning in the past can be reasonably well reconstructed fromrecords of" the amounts of coal and ligni te rnined and of theamounts of oil and natural gas taken from the world t s "wells".
From a knowledge of the mean carbon content of these fuels we canconvert the tons of coal and lignite, the barrels of oil, and the
cubic feet of natural gas to moles of C02 gas.
A summary of theamounts of CO 2 produced each year from fossil fuels is shown infigure 10-1.
The cumulati ve amount of CO 2 produced up to any
gi ven time is also shown in this figure.
As of 1980 this total
was about 14 x 10 15 moles.
For comparison the atmosphere in 1980
contained 55 x 10 15 moles of C02.
In addi tion to the CO 2 produced by the burning of foss ilfuels, man has added CO 2 to the atmosphere through his fores try

500

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1900

1920

YEAR
Figure 10-1. Plots of the annual production of fossil fuel C02(upper) and the integrated production of fossil fuel C02 (lower)
from 1860 to 19BO.
The relative contributions of coal andlignite, of oil and of natural gas are indicated in the upper
panel (559,562).

501

and agricul tural practices.
About 10% of our planet's land has
been harnessed for agriculture. Some of this land was originallyforested.
Tilling increases the oxidation rate of the organic
matter contained in soils and thereby decreases the global inventory of these compounds. We have also cut forests which are stillin the early stages of regrowth.
The mass of carbon present intoday's living material is about equal to the amount of C02 in today's atmosphere.
The mass of carbon in soils is 2 to 4 times
that in today's atmosphere.
While a decrease in the mass of living material and in the mass of carbon stored as organic compounds
in soils has contributed to the rise in atmospheric CO 2 content,
our records of these activities are not nearly good enough to permit the exact accounting possible for fossil fuels.
Rough estimates suggest that over the last 200 years soils and forests havecontributed an amount of C02 to the atmosphere comparable to thatfrom fossil fuels. A scenario for these releases is shown in figure 10-2. As it is obtained by an indirect method to be discussed
later in the chapter, we shall not say any more about it at thispoint.
Before moving on to the future, it is worth mentioning a few
of the imponderables involved in any attempt to evaluate thesetrends in terrestrial biomass.
One of these imponderables has to do wi th the regrowth ofpreviously cut forest.
An idealized response function shown infigure 10-3 demonstrates the problem. First of all, when a forest
is cut the biomass does not disappear immediately.
The wood may
be used in bUildings or for making paper.
The roots (which constitute up to 50% of the forest biomass) are left in the ground to
decay.
The small branches and bark are often left to be eaten bybacteria. Where material is burned on the site, ten or so percentof the carbon is converted to charcoal and becomes a semipermanent consti tuent of the soil.
Thus the biomass associatedwith a cut forest does not immediately disappear.
Rather, like a
mixture of radioacti ve isotopes wi th widely varying half-Ii ves,
its disappearance has a long history.
Soon after a forest is cut new growth begins.
If uninterrupted, this growth will on the time scale of a few tens to a few
hundreds of years produce a new steady state biomass. More likely
the forest will be cut again when the timber company or Swidden*
farmer judges it "ready".
Thus many of the forests of the world
are somewhere in the midst of a regrowth.
Of these some are in a
decay phase where the decomposition of the cut material outweighs
new growth..
Some are in the growth phase where the formation of*In developing countries where commercial fertilizer is too
expensive farmers often rotate their lands.
The farmer cultivates a plot of land for several years and then when its soil
nutrients become depleted he allows it to go "wild" and moves toa new plot.
Forest growth in the wild patch renews its nutrients (presumably by drawing on nutrients from deeper in thesoil, through nitrogen fixation, and via atmospheric fallout of
P04, K, S, ... ).
A decade or so later the farmer returns thisplot to cultivation, completing the cycle of his acreage.
This
type of rotating land use is called Swidden farming.

502

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INTEGRATED PRODUCTION
15AS OF 1980 (10 mol)
FOSSIL FUEL C02

14

FOREST a SOIL C02

19

TOTAL

33
TOTAL

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1880

1930

FUEL CO

2....
1980

YEAR
Figure 10-2. A scenario (based on the tree-ring-based atmospheric
13 C/12 C record) for the net release of CO 2 from the forests andsoils over the last 200 years.
Also shown is the documented
record of fossil fuel C02 release.
The total production fromthese two sources is also given.

503

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TIME AFTER CUTTING (yr)Figure10-3. Hypothetical curves showing the change in carbon
storage experienced by a forest plot after cutting.
The curvelabeled new biomass represents new forest growth.
The curve
labled old biomass represents the composite decay curve forleaves, bark, roots, and wood (or wood products) present on theplot at the time of cutting.

504

new timber outweighs decomposition.
The point of all this is that
it is the differential rather than absolute rate of forest cuttingthat matters.
To assess forest biomass changes from estimates offorest cutting alone would be equivalent to judging a man's financial status by looking only at his expenditures!
We have already mentioned a second imponderable.
When forests burn some of the carbon goes to the iner't form of charcoal.
By contrast, when tr'ees experience "natural" deaths, bacteria (andtermi tes) do a mOr'e complete job of recycling the carbon.
In theindustrial countries an attempt has been made to control forest
fires.
These containment efforts have been rather successful.
Forests in these areas are destroyed less frequently by fire thanprior to man's "management."
In developing countries both grass
and forested areas are frequently burned.
Presumably this practice has intensified along with population growth. Thus, while itis hard to tell how much charcoal pr'oduction has enhanced the carbon inventory in soil, this contr'ibution is unlikely to have been
negligible (571).
A third imponderable has to do with the anthropogenic fertilization of forests.
Automobile engines and fossil-fuel-powered
electr'icity generators fix atmospheric N2 (i.e., make oxides ofni trogen) •These oxides go into the air and ar'e eventually conver'ted to N03, which is carried to the earth (and sea) by rainfall
and aerosol impact.
Some of the NH 4 0H used to fertilize agricultural lands is released to the air as NH 3 gas.
When this NH3 encounters H 2 S0 4 - or HN03-bearing aerosols it combines as part of
the aerosol and soon returns to the earth's surface.
Rains alsodeliver appreciable P0 4 to the land surface.
While the origin of
this phosphate is not well understood, part of it is anthr'opogenic(for example, from coal burning).
The C02 content of the air isalso steadily rising.
Thus, three of the limiting ingredients forterrestrial plant growth are being made available through man's
acti vi ties.
A weal th of expe rimental da ta shows that gi ven more
nutrients plants grow faster. Presumably the trees in our forests
are, on the average, growing a bit faster' as the r'esult of thisinadver'tent fer'tilization.
Presumably these higher' growth r'ates
would lead to a greater forest biomass than would exist in theabsence of these fertilizations.
As of 1980 the CD2 content of
the atmosphere was about 20% higher than prior to the turn of the
century.
It is surely safe to place this percentage as a firm
upper limi t on the incr'ease in fores t photosynthesis r'ates.
The
true growth enhancement is likely considerably smaller'.
The ;list of imponderables does not end here.
However, thepoint has been made.
When considering the terrestrial biomass weare dealing with a complex and very heterogeneous system. Our information base is simply inadequate to yield the kind of answer's
we would like.
Hence, we will have to find ways to work ar'ound
the r'ather large uncertainties in the time history of the net CO 2
release by the terr'estrial biosphere and soils.
CO 2 PRODUCTION IN THE FUTURE

First,

So much for the past.
What can we expect in the future?
we need to know the magnitude of our reserves.
They are505

Table 10-1. Estimate of recoverable fossil fuel16 reserves. For comparison today's atmosphere contains 5.5 x 10
moles of C02 and16
today's fores ts contain about 5.5 x 10
moles of carbon i 6 The
amount of fossil fuel C02 produced as of 1980 was 1.4 x 10
mole
(560,561).

Fuel

Mass of Carbon
10 16 moles

Coal and lignite
Petroleum
Natural gas
Oil shale
Tar sands
TOTAL

~

30
1.9
1.2

1.4
0.6
- 35

Table 10-2. C02 production per unit of energy produced (560).

Fuel

MethanePetroleum
Coal
Oil shale
Tar sand
Synthetic methane
Synthetic petroleum

10- 6 Moles CO2
calorie energy

5.0
6.6
8.3
. . 13-13. . 12. . 12* Relative to methane.

506

C02 Production
Index *

( 1)
1.3
1.7
2.6
2.6
2.5
2.5

listed in table 10-1.
Together they add up to a total of six
times more carbon than is now in the air!
The greater part ofthis is coal.
Thus the future C02 levels in the air depend verymuch on the extent to which we utilize these coal reserves.
However, it must be kept in mind that even if we were to burn only
10% of this coal, and a major portion of the oil and gas, we would
still generate an amount of CO 2 equal to that presently in the atmosphere.
To provide some idea as to how fast we might run
through our fossil fuel reserves, if the present rate of fossil
fuel use were maintained, we would generate an amount of C02 equal
to that in the atmosphere in about 65 years.
One has only to consider the events of the 1970's to realize
how hazardous it is to make long-range energy-use forecasts.
A
typical forecast of late 1960 called for growth rates of from 3%
to 4% per year in energy usage.
Then came the Arab oil boycott.
Soon after came the rapid increases in oil prices. As can be seenin figure 10-4 the growth ra te dropped drama tically.
Seriousquestions with regard to the safety of nuclear power plants andconcerns wi th regard to the storage of the radioacti ve "ashes"from these plants have been raised during the last decade. Reacting to these unforeseen events energy growth predictions have been
dramatically modified.
The example shown in figure 10-5 is that
released by Exxon at the time this book was being written.
Itcalls for a 2% per year growth rate during the next 20 years.
Even as you read this book the situation is likely to have changed
again.
This is why we say that any model of future C02 levels
must accommodate itself to a wide range of scenarios.
While we have no intention of making our own prediction, a
few paragraphs about the ingredients of such scenarios is inorder. The amount of energy used must be related to the number ofpeople who inhabit our planet.
Currently the population of the
globe rises by 2% per year.
At this rate the population will
double in -35 years.
While this growth ra te is gradually being
stemmed by more effective birth control practices few believe thatthe population will level off before reaching 8 billion (asof1980 it was about 4.5
billion).
More pessimistic outlooks would
predict growth to continue until the population is 12 to 16 billion.
As the standard of Ii ving of the world's poor people isclosely related to their per capi ta energy consumption, despite
energy conservation efforts by the world's "rich", the demand forenergy will certainly rise along with population.
In this connection it must be remembered that 3 out of 4 people are in the"poor" category and that the growth in population is almost entirely through children born to "poor" families.
Eventually it
will be the energy habits of the "poor" rather than of the "rich",which will determine global energy use.
This rapidly expanding
demand by the people of the developing countries is reflected in
Exxon's prediction of the distribution of world energy demand (seefigure 10-6).
In order to make a prediction of future C02 production rates,
we also must estimate what fraction of our energy will come fromfossil fuels.
During the 1970' s about 90% came from this source(see figure 10-7).
The prospects that this situation will change
dramatically in the next several decades are indeed small.

507

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100

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1950 1955 1960 1965 1970 /975

1980

YEAR
Figure 10-4. Logari thmic plot of CO 2 emissions from fossil fueluse as a function of time from 1950 to 1980 (upper panel).
The
slopes of the line segments on such a plot are measures of thegrowth rate of these emissions (shown in % per yr).
Note thesharp drop in growth rate which took place in 1973.
The lower
panels show separate plots for coal, oil, and natural gas.
Bothoil and natural gas usage unde rwent a ve ry sharp drop in growthrate in 1973. Coal usage did not show this change. This compilation is the work of Ralph Ratty of Oak Ridge Associated Uni versities (562).

508

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E

v
--0

-....

z
o

I-

u

:J

o

o

a::

0..

...J

«

:::>

z
z

<t
1970

2010

1990

2030

2050

YEAR
Figure 10-5. Plot showing projections of fossil fuel use.
The 3%
per year growth curve represents a median estimate made prior to
the Arab oil boycott and the OPEC price hikes.
The Exxon projection was made in 1980 (691).
The squares represent historic
production records.

509

>«
o

250....---.-----.. .

-----....-----~

0::

W

cQ.
Z ..J t-

11~
«-Z
~OW

WIL.--100«>- (J)> 125....J_

C)

0:: W ::>

W~o

100

W m

z «w

75

z
o

50~~==~========~===::::J,

....J

25

...J
~

O~-""-----""------L.-----~
1970
1980
1990
2000

YEAR

_

Figure 10-6. Historic (up to 198-0) and projected (1980 to 2000)
energy use for various countries and blocs of countries.
The area
designated "other" is made up mainly of the developing countries.
The projections were made by Exxon C-Orporation
1981 (691).
~
250.-----...-----.....-----....--.....;;..----.

©

>«
o

225

>-n:
-1
W

200

&a.

::l..Jt- 115(/)OZ->- IL. w 150
C)O..J
a::: (/) «W....I>ZW-

W 0::::>
00:: 0
..J
W

«
o::m

I25

100
20%

15
2%

Oz
3:0
--I
--I
~

47%
42%

38%

-----...L.-----."
1990
2000

O"'--~~---.......

1970

1980

YEAR
Figure 10-7. Historic (up to 1980) and projected (1980 to 2000)
estimates of the energy supplied from various sources. As of 1980
fossil fuels accounted for 92% of the total.
The projection madeby Exxon C.orporation
1981 is that, as of the year 2000, fossil
fuels will still account for 82% of the total (691).

©

510

Resistance to the large scale usage of nuclear power runs high.
Despite the considerable talk of solar power, its fraction of the"pie" will remain small until a major breakthrough in technology
comes which will permit electrical energy to be produced at acompeti ti ve price.
In a world pushed to grow enough food andfiber it is difficult to believe that vegetation will become amajor source of energy.
And so it goes; the end to the dominanceof fossil fuels is surely many decades away.
Another ingredient to C02 production estimates comes fromthe fact that per unit of energy delivered to the user quite
different amounts of CO 2 are delivered to the atmosphere dependingon the kind of fuel used (see table 10-2). Methane, which has thelowest C02 production to energy delivery ratio, is used as areference.
For oil the ratio of C02 produced to energy deliveredis 1.3 times greater, for coal it is 1.7 times greater and forsynthetic fuels (methane and liquid hydrocarbons manufactured fromcoal) and for oil derived from tar sands and shales the ratio
averages 2.5. Thus we see that, as we run out of natural gas andoil and begin to use other fossil fuel sources, the amount of CD2
produced per unit energy consumed will rise dramatically.
Finally, it should be mentioned that as the population increases and as fuel prices rise the temptation to even further reduce the planet I s standing forest stocks will dramatically rise.
Forests will be converted to farms and forests will be cut forfuel.
We currently have more carbon in our forests than in ouroil plus natural gas reserves.
CAPACITY OF THE SEA FOR FOSSIL FUEL C02 UPTAKE

The distribution of the C02 produced by man's activities between the air and the sea is a function of both thermodynamic(i.e., capacity) and kinetic (i.e., rate) factors.
\oJe will consider the _capacity factors first.
Although C02, like any othergas, dissolves in the sea, were simple solution the only mechanism
for its uptake then the C02 produced by man would remain almost
entirely airborne.
The ocean would take up only about 3% of that
added to the air.
Significant amounts of C02 enter the ocean only
because of the reactions:
2HC03andH 20 + C02 + CaC0 3

10-1- Ca++ + 2HC03- 10-2~

The C03= in reaction 10-1 is that which is dissolved in theocean. The CaC03 in reaction 10-2 is that in the sediments on thesea floor.
If present in great enough quanti ties these naturalbases will eventually neutralize the man-made carbonic acid.
Is there enough base to neutralize the amount of C02 which
would be produced if all our fossil fuel reserves were burned?
The sea contains a total of about 12 x 10 16 moles of CO 3= iOlf6
The upper 10 cm of marine sediments* contain about 8 x 10
*We take 10 em as this is the depth to which burrowing organisms
penetrate.

511

moles of CaC03. Our combined fossil fuel reserves amount to about
35 x 10 16 moles of carbon (see table 10-1).
Thus viewed in thisway the system could neutralize only about one-half of the carbonic acid man might ultimately produce.
As it turns out~ both the estimate of the amount of base in
the sea and that in the dynamically available sediment are toolow.
Our use of the mean amount of CO 3= ion in the sea as ameasure of its capacity to take up C02 is misleading.
Becausemost of the water in the ocean is deep water, this inventory isheavily biased toward the carbonate ion content of deep water.
Actually, the atmosphere "sees" only the carbonate ion in surfacewater ~ which, as shown in chapter 3, ranges from 1.3 times themean deep water value in regions of cold surface water to as muchas three times the deep water value in regions of warm surface water.
Hence~ the capacity of the sea to neutralize C02 is larger
than the value of 12 x 10 16 moles gi ven above.
As most of the
deep water in the ocean is cold it is the C03= ion of cold
surface water which is most important.
Hence we should increase
our ocean capacity estimate by about 25 percent to a value of 15 x
10 16
moles.
This
is
a
more
realistic
estimateoftheneutralization capacity of the sea itself.
Our estimate of the neutralization capacity of marine sediments is also low.
We assumed that only that amount of CaC03 lying within 10 em of the sediment-water interface is available forsolution on a reasonable time scale (i.e., several thousand
years)~
the reasoning being that molecular diffusion throughseveral centimeters of sediment coupled with the low solubility ofcalcite would yield extremely low solution rates. Hence, we focus
on the layer stirred by organisms.
This mixing would continuallybring new calci te close to the sediment wa ter interface removing
the necessity for a long diffusion path. This reasoning is in error in that as the calcite in the upper 10 em of sediment is dissolved, organisms would reach down into sediment which had previously been beyond thei r burrowing range.
Indeed calci te would
continue to reach the interface as a consequence of their churnings until a layer of calcite-free residue 10 em thick had been
generated!
The amount of calcite kinetically available thenbecomes:x 10 g/cm 2

M=

10-3

1 - fCaC03
where
ment.

fCaC'0 3 isthefractionofcalciteintheoriginalsedi-

In order to obtain a rough estimate of this capacity marine
sediments can be divided into two categories:
calcite oozes
covering about 20% of the sea floor and averaging 85% calci te bymass and red clays covering about 80% of the sea floor and containing negligible calcite.
The dynamically available calcite in
the carbonate oozes would then be:

0.85
M

x 10 = 57 g/cm 20.15

0.57 moles/cm 2

512

The total calculated in this way comes out to be about 40 x 10 16
moles.
These revised estimates are more promising!
The total neutralization capacity is half again the amount of carbon locked up
in recoverable fossil fuels.
The bulk of this capacity, however,resides in the sediments rather than in the water column.
As weshall see, it will takeman's C02 several thousands of years to
get at this CaC0 3 .
UTILIZABLE CAPACITY - SIMPLIFIED CALCULATION
While the approach we have just taken provides a ready meansof getting at the til timate buffering capaci ty of the ocean, it
says nothing about the amount of this capacity utilizable for any
given amount of CO 2 added to the atmosphere.
As we are involved
wi th a reversible chemical reaction, only a fraction of the CO 2
will be neutralized before an equilibrium distribution is achieved.
Once the pressure of C02 in the atmosphere has fallen to theequilibrium pressure, the neutralization will proceed no further.
In the discussions which follow, we will consider separately
the cases of neutralization by the carbonate ion in the sea alone
and of the sea plus the calcium carbonate in its sediment because
the time required for atmosphere and sea to come to equilibrium ismany hundreds of years while the time for atmosphere, sea, andsediment to come to equilibrium is many thousands of years.
Thecase for the sea alone will be considered in this section.
We will first consider the situation for an ocean with uniform chemistry and with no borate.
This simplification is usefulin that it allows us to understand the big picture without mathematical complexity.
The equilibrium equation for reaction 10-1 isas follows:[ HC0 3-]2
10- 4
Ka =
[ C0 3=] PC0 2
where K is the equilibrium constant for reaction 10-1 (see table3-7) and a is the solubility of C02 in sea water (see table 3-7).
In our hypothetical borate-free ocean we can write down thefollowing equation relating alkalinity to the concentrations of
CO 3= and HCO 3-:[Alk] = [HC03-] + 2[C03=]
10-5
Eliminatin~

the[HCO 3-]concentration

betweenthetwoequationsyields:Ka =([Alk] -2[C03=])2

PC02 [ C0 3 =]10-6This equation contains only [C03=] as an unknown.
A rough es tima te of what happens \'1hen addi tional CO 2 isadded to our simplified ocean-atmosphere system can be obtained byassuming that the numerator of this expression remains constant.

513

This approximation yields the following relationship:
10-7

PC02 [ C0 3=J = pOC0 2 [ C0 3=Jo

Thus to the first approximation for each 10% the C02 content of
the atmosphere rises the C03= ion content of surface water dropsby 10%.
C03= ion constitutes between 7% (for cold water) and
12% (for warm water) of the LC02 in the surface ocean.
As the total CO 2 content rises by the same amount as the CO 3= ion content
falls (see reaction 10-1) a 10% rise in PCOz will lead to a 0.7%
rise in LC02 for cold surface water and a 1.2% rise in LC02 forwarm surface water.
What we seek is the equilibrium fraction of C02 remaining
in the atmosphere.
In honor of Charles D. Keeling whose monitoring of the atmospheric C02 forms the foundation for all studies ofthe fate of fossil fuel C02 we sill refer to this fraction as the
Keeling fraction, fK, where:

M*atm

10-8

M*atm + M*ocean
The amount of fossil fuel C02
is given can be approximated by:accumulated

in

the

air,
10-9

M*atm
pOC02

where pOCO 2 is the partial pressure of CO 2 in the air prior to
the industrial revolution and MO atm is the mass of C02 in thepreindust~ial atmosphere.
The amount of excess C02 taken up by
the ocean can be approximated by:where [C03=Jo is the mean carbonate ion content of the surface
ocean prio~ to the industrial revolution and Vocean is the volume of the ocean.
It is the drop in CO 3= ion concentration insurface water that provides the measure of the amount of C02 taken
into the sea (one mole of C03= is utilized for each mole of C02
removed from the atmosphere).
Thus fK is gi yen by the following
equation:
M° atm10-11

514

Using the approximate relationshipabove, this equation becomes:

betweenPcD2 and

C03=

given

1
10-12

The values of the constants in this equation are as follows:
20
Vocean = 13.3 x 10
liters[C03=Jo = 200 X 10- 6 moles/liter
MOatm = 5 x 10 16 moles
pOC0 = 2~0 X 10- 6 atm
2
Thus:
1

10-13
1 +

5.3

(1 -

280/ PC0 2)(PC0 2 /280 - 1)
It can be seen that for small increases in atmospheric C02 content
the Keeling fraction lies close to 1/6.3 or to 0.16.
On the other
hand, as the C02 content of the atmosphere becomes very large, theKeeling fraction approaches unity (the CO 3= ion content of theocean would eventually reach very low values and the sea wouldcease to take up the excess CO 2 added to the atmosphere).
The relationship between the Keeling fraction and the atmospheric C02
partial pressure given by this simplified equation is shown infigure 10-8.
The total amount of fossil fuel C02 generated is related to
Keeling fraction and the atmospheric C02 partial pressure as
follows:

10-14
For the simplified assumption that the product of surface ocean
CD 3= ion concentration and atmospheric CO 2 partial pressure remains constant the relationship between MFF and PCO
z is shownin figure 10-8.
It can be seen from this figure that In order toraise the atmosfhere's C02 pressure by a factor of two (i.e., toabout 600 x 10- atm) an amount of C02 equal to about 4 times thatin the preindustrial atmosphere would have to be released by fossil fuel C02 combustion.
It must be remembered that this calculation is made with theassumption that the entire ocean (but not of the sediments) isthoroughly equilibrated with the atmosphere.
As we shall see overthe next century or so, only a rather small fraction of the ocean
will equilibrate with the air.
Thus the Keeling fractions given

515

z

.40

0

-~

[CO;] x Peo.,.,..

=CONST.

2~,.",,APPROXIMATJON,,""'-

U

« .30
a:

"..

,.",,-

"..

LL

C)

z .20
..J
lJJ
W
~

.1°300500

700

P~~; (I0-6 atm )

30

6

....E

o 0

-

//~-~!

4~[C0 3] x PC02 = CONST•......
'to'to-~900

APPROXIMATION

2...__......

~-_......,'+-

10 ~'to-....__....__...........__.....__.... o900

Figure 10-8. Plot of the Keeling fraction as (upper) and of theamount of fossil fuel C02 added to the atmosphere (lower) as afunction of atmospheric PGD2' assuming that the atmosphere andentire ocean maintain equillbrium, and that no interaction wi th
sedimentary GaG03 occurs
Two sets of curves are gi ven.
One isfor the simplified si tua tion where the product of CO 3= concentration in surface water and the PG02 partial pressure in theatmosphere is assumed to remain constant.
The other ~s based on
rigorous calculations outlined in the text.
e516

in figure 10-8 are lower limits.
As our interest is centered on the CO 2 buildup in the coming
century, it is appropriate to repeat these calculations for thecase where only 10% of the ocean volume is equilibrated with theatmosphere (as we shall sec, the effective equilibration volume
during the next 100 years will be in the range of 10% to 20% of
the ocean volume).
The Keeling fraction for equilibration between
the atmosphere and 10% of the sea is as follows:
1

10-15

The graphs in figure 10- 9 show the rela t'ionships between fK and
PC02 and Mfi'F and PCO
obtained in this way.
As can be seenfrom this figure, the ~eeling fraction rises from 0.65 for small
amounts of CO 2 production and approaches uni ty for large amountsof C02 production.
In this case the amount of C02 production required to yield a CO 2 pressure of 600 x 10- 6 atm is only about 1.5
times the amount of CO 2 in the preindustrial atmosphere.

UTILIZABLE CAPACITY - RIGOROUS CALCULATIONLet us now do the ocean-atmosphere problem in a more rigorous
way. There are seven equations governing the steady state distribution. As before, we have the eqUilibrium for reaction 10-1:
aK :::[HC0 3-J 2

10-16[CO 3 = ] PC0 2

Next we have the full alkalinity equation:
10-17

Further wereaction:

have

the

equilibrium

equationforthechemicalwhich is:10-18

We also have the mass balance equation for the dissolved inorganiccarbon species:
10-19

and the mass balance equation for borate species:517

z 0.90 .....- -......---...--...,..--....,.--.....--...,o.....
[CO;] x PC02 = CONST.
u
0.80 APPROXIMATION ~
«
a::
L&-

__ -

-

- -

--

"FULL EQUATIONS

C)

Z
-l
W
lJJ
~

EQUILIBRATION WITH
10% OF OCEAN VOLUME
0.60

300

500

700

900

p~~~ (IO-6 atm )E

3 .....- - - . - - - - . . . - -......- -......- -......---..15
FULL EQUATIONSEQUILIBRATIONo"\ .,,-~
2 WITH 10% OF
10 E
toOCEAN VOLUME

7

o

-

[cO~]XPC02 = CONST.
APPROXIMATION

0 ....- - - . . . - -....- -......----------------....... 0

300

500

700

900

p ~6~ (I0-6atm )Figure 10-9. Plots of the Keeling fraction (upper) and mass offossil fuel C02 added to the atmosphere (lower) versus atmospheric
PC02' assuming that only 10% of the ocean volume equilibrates
witn the atmosphere (and, of course, that no interaction withsedimentary CaC03 occurs).
The dashed line is for the simplified
calculation where the constancy of the product of the surface
water carbonate ion concentration and the atmospheric C02 partialpressure in the atmosphere is assumed.
The solid line is for the
rigorous calculation outlined in the text.

518

10-20
Finally we have the equation for the Keeling fraction:(pco 2 /poC0 2 -1) MOatm

10-21

and the equation starting the relationship between the massfossil fuel COz produced and the Keeling fraction:

of

10-22

These seven equations contain seven unknowns([H zBO 3-] ,[H3 B0 3],[reOz], [HC03-], [C03=], fK and MFF).
~hevalueso~[LCOz]O,
pOCO z '[Alk],
[LB],
a,
K,
KB,
M atm,
andM ocean are knowns([Alk] does not change in the absence ofinteractions with the ocean sediments).
Were the ocean homogeneous in its properties we could apply
these equations as we did the simplified ones gi ven above.
However, as we are now trying to get the best answer possible fromthis approach, we must take the spatial distribution of properties
in real ocean into account.
Our strategy will be to divide theinterior of the ocean into a number of subunits and to establish
for each the properties for its source waters.
The subdivision will be made according to temperature.
Boundaries are placed at intervals at 4°c (i.e., at 4°,8°,12°,16°, 20° and 24°C).
The volumes of ocean water in each of thesesubdivisions are listed in table 10-3.
We also need to establish
the C03= ion content of surface waters as a function of temperature.
The C03= ion concentrations for all the surface water
samples taken as part of the GEOSECS program are plotted against
temperature in figure 10-10.
A strong correlation is found; thewarmer the water, the higher its C03= concentration.
The trend
seen in this diagram is approximately that expected from the temperature dependence of the equilibrium constants for surface waters
equilibrated with the COz partial pressure of the atmosphere.
This can be seen from the dashed line shown in the figure which
depicts the carbonate ion content of a single water sample with analkalinity of 2330 ~eq/kg equilibrated with the atmosphere (peoz6= 330 x 10atm) at temper,atures ranging from a to 30 0 e. The departures from this trend reflect the fact that alkalinity, surface
water salinity and surface water peoz vary from place to place.
While the carbonate ion content of a given surface water provides a good measure of its capaci ty to take up excess eo z fromthe atmosphere, there is another parameter which proves more precise in this regard.
It is called the Revelle factor, R.
It is
the ratio of the fractional rise in atmospheric COz content to thefractional rise in the Leo z content of a water sample at equilibrium with the atmosphere, i.e.,

519

Table 10-3. Subdivision of the ocean for the purpose of calculating the equilibrium responseof the ocean to a doubling of atmospheric CO 2 content.
Temp. Volume
Range

Fraction
Total
Ocean
VolumeOutcrop Outcrop Outcrop Revelle
Factor+
rC02+ C03=+
Alk(1850) (1850)(1850)
llm/kglleq/kg
llm/kg

Outcrop Outcrop Revelle
Mass**
C03=*
Factor* Excess CO2
rC02*(future) (future) (future)(fu~ure)llm/kglJffi/kg
10 1 moles°C

10 16 m 3

>24

1.0

0.008

2330

1938

276

8.5

2097

172

11.1

0.16

20-24

0.7

0.005

2330

1975

249

9.0

2127

151

11.9

0.11

c..n

16-20

1.7

0.013

2330

2012

223

9.5

2156

132

12.9

0.25

o

12-16

2.7

0.020

2330

2048

197

10.2

2183

114

14.0

0.37

8-12

3.3

0.025

2330

2081

173

11.0

2208

98

15.2

0.43

4-8

14.9

0.113

2330

2112

152

11.8

2231

84

16.3

1.82

<4

107.7

0.816

2330

2139

135

12.7

2252

73

17.2

12.43

rrOTAL 132.0

1.000

-

-

-

-

-

-

-

15.57

I'..)

+For atmospheric CO 2 pressure equal to 280 x 10- 6 atm.
*For atmospheric CO 2 pressure equal to 600 x 10- 6 atm.
**Equi1ibrlum mass of CO 2 taken up by the particular ocean subvolume for an atmosphericpe02 of 600 x 10- 6 atm. For comparison the amount of excess C02 in the atmosphere forthis pe02 is 11 x 10 6 moles.
.

275

250

..........

225

0'

• ATLANTICA INDIAN
o PACIFIC

0

eX.

-"-0

E 200
:l

'-'

1Itt')

a

175

0

u
150

5

10

15

20

25

30

TE MPE RATURE (OC)
Figure 10-10. Carbonate ion concentration measured in surface
waters of the ocean during the GEOSECS program as a function of
surface water temperature.
The dashed line represents the values
expected if a single water sample were equilibrated with an atmospheric CO 2 pressure of 330 x 10- 6 atm at different temperatures.
The data points are based on measurements made during the GEOSECS
program (406).

521

6PC0 2 /PC0 2

R

10-23

6EC02/EC02
or:
R

10-24

A rough idea regarding the relationship between the Revelle factor
to the carbonate ion content can be obtained using the approximations that:
10-25
and that:
10-26
If equation 10-26 is divided by [EC02Jo and if equation 10-25
used to eliminate [C03=J from equation 10-26, we obtain:is6[EC 0 2J ~[PC02 - POC0 2] [ C0 3=Jo
[ EC0 2]O
Rearranging
thispOC02 we obtain:

equation

6[ EC0 2][ EC0 2]O

PC0 2
and

substituting

10-27
6PC02

for

PC0 2 -

[C03=Jo
10-28

Hence:
10-29
While this expression gives the approximate value of R in order toobtain its precise value, we must use the set of equations given
above.
We can now check the extent to which this approximation
holds.
Shown in figure 10-11 are the Revelle factors for surface
waters from throughout the Atlantic Ocean plotted against the EC02
to C03= concentration ratio.
For warm surface waters
theEC02/C03= approximation gives a value which is about 10% too
low.
For cold surface waters the approximation gives a value
which is about 25% too high.
It must be pointed out that just as
the [EC02]/[C03=] ratio for any given surface water sample' increases with the atmospheric CD2 content so also does the Revellefactor.
In carrying out calculations for progressively increasing
atmospheric C02 contents, the Revelle factor must be recalculated
as the C03= ion content of surface ocean water falls.
Now back to our rigorous static calculation.
Having established the volumes and outcrop properties of the various subvolumes of the ocean (see table 10-3) we can apply the full set of

522

20
18

COLD
SURFACE
WATERS

•
I

16
Uri)

0

~

14

C\J

0

u

12

V'I
10

8
WARM SURFACE WATERS

6 8

10

12

14

16

18

20

REVE LLE FACTOR
Figure 10-11. Comparison of the ratio of ECOz to C03= concentration for surface waters with the Revelle factor calculated forthese waters.
The Revelle factor is the ratio between the fractional change in the COz partial pressure and the fractional change
in ECOz content for the water.

523

equations given above to each.
The results for one value of
peo 2 (600 x 10- 6 atm) are listed in table 10-3.
Plots for fK
vs. pe02 and ~F vs. peo 2 are given in figure 10-8.
The
results are not much different from those obtained using thesimple approximation.
Similar rigorous static calculations for an equilibration
with 10% of the ocean volume yields the plots of fK vs. Pe0 2
and of MFF vs · Pea 2 gi ven in figure 10-9.
Again they do notdiffer greatly from those obtained when the simplifying assumptions were used.
KINETICS OF FOSSIL FUEL C02 UPTAKE BY THE SEA

While the calculations leading to the partition estimates infigure 10-9 handle the capacity factors for the sea correctly the
kinetic factors are handled in a highly simplified manner. It was
assumed that 10% of the ocean fully equilibrated with the atmosphere and that the remaining 90% was totally isolated.
In
reality the ocean would be expected to show a gradation from nearly complete equilibration for surface waters to no equilibration
for the older deep waters.
If we are to get a meaningful answer
then a model must be developed which incorporates the kinetic resistances to equilibration (namely the finite rate of transfer of
CO 2 across the air-sea interface and the finite rate of mixing between surface and interior waters).
Before doing a proper kinetic model let us consider theorigin of the 10% figure used for the calculations summarized infigure 10-9. It is based on the average of penetration depths fortri tium as measured at the stations in the Atlantic and Pacific
Oceans occupied during the GEOSECS program (see table 10-4).
As
the mid-date of the addition of tritium to the ocean was 1963 andthe mid-date of the Atlantic and Pacific GEOSECS surveys was 1973,
this average penetration depth gives an estimate of the fractionof the ocean contaminated with tritium on a ten-year time scale.
As shown in table 10-4, the result comes out to be 375 m(corresponding to 10% of the ocean's volume).
In applying the tritium-based penetration depth we must firstconsider the difference in the time available for the averagefossil fuel C02 molecule to penetrate into the sea and the time
the average bomb tritium atom had to penetrate into the sea (as of1973). At first glance one might be tempted to take the ratio of
the mean-age of fossil fuel C02 molecules (i.e., 30 years) to themean-age tritium atoms as of the time of the GEOSECS survey (i.e.,
10 years). This would be incorrect because the C02 age reflects
the mean time after addition to the atmosphere, while the tritium
age is the time since entry into the sea.
The nume ri cal mode I
developed in the section which follows shows that the equivalentocean residence time for fossil fuel C02 molecules is 0.57 theirmean age.
Thus the ocean residence time for fossil fuel CO 2
molecules is 17 years.
The longer a substance resides in the ocean, the greater will
be its penetration depth.
While we do not know exactly what thetime dependence is, our observations of bomb tritium show that thevolume does not grow linearly with time (i.e., after 30 years have524

Table 10-4. Summary. of the mean tri tium penetration depths as of 1973 for various partsof the ocean.
The numbers in the right hand column are the product of the mean penetration depth for a gi ven zone times the fraction of ocean area consti tuted by that zone.
The penetration depths for the Indian Ocean are assumed to be the same as the average for
the Atlantic and Pacific.
Based on tritium measurements made by Ostlund at the University of Miami as part of the GE08ECS program (21,22,23).

(.J"'1r-..:>

Atl.
rn

Pac.
m

Global Meanm

Frac. Ocean
Area

C:ontribu tion to
Global Mean Depthm

Antarctic (>45°8)

270

475

375

0.169

62

s. Temp.

(45°8-15°8)

340

340

340

0.264

90

Equator (15°8-15°N)

220

225

225

0.281

63

N. Temp. (15 ° N- 45° N)

700

490

560

0.192

107

N. Pac. (>45°N)

-

310

-

0.028

9

N. Atl. (>45°N)

1060

-

-

0.036

38

Arctic (->70 0 N)

-

-

-200

0.032

6

(.J"'1

Global Mean

375

passed, tritium will not have reached three times the penetration
depth observed at the time of the GEOSECS surveys).
A more appropriate estimate appears to be that the penetration depth increases as the square root of time (as would be expected if itwere being moved downward by eddy diffusion).
If so then after 17
years the mean depth would be 1.3 times that for 10 years.
Because of the finite rate of C02 exchange between the surface ocean and atmosphere, even the uppermost waters of the ocean
are not quite at equilibrium with the C02 content of the air.
Thenumerical to be presented below suggests that at present it has
only o. e6 the equilibrium amount of fossil fuel C02.
Hence thetritium-based mean penetration depth must be multiplied by 1.30 toaccount for the longer time available for C02 penetration and by
0.86 to account for the failure of the surface waters to achieve
full equilibrium with the atmosphere.
The result is 420 meters or
11% of the ocean volume.
This type of analysis can be extended to the future.
For any
fuel use scenario we can compute the mean age of the fossil fuelCD2 molecules present in the atmosphere ocean system at any timein the future.
In figure 10-12 are given several fossil fuelscenarios and their corresponding mean C02 ages.
If we confine
our attention to the next 80 years, these graphs make it clearthat we are interested mainly in those portions of the ocean which
are ventilated by fossil fuel C02 molecules with mean ages in therange 30 and 60 years (i.e., effective ocean penetration times in
the range 17 to 34 years).
Figure 10-13 shows a possible relationship between the fraction of ocean ventilated and the square root of the penetrationtime.
A straight line joins the origin and the bomb tritium calibration point (i.e., 10% ventilated in 10 years).
The range ofventilation times consistent with the distribution of naturalradiocarbon in the sea is shown to straddle the extension of thisline lending support to the approximation that the volume equilibrated increases as the square root of penetration time.
The
range for the natural calibration radiocarbon stems from the difference between the deep sea ventilation time for radiocarbon (1000 years) and the ventilation time for heat and helium (- 400
years) .
Following the arguments presented in chapter 3, thedeep-sea ventilation time for C02 should lie somewhere in between
those for 14C and heat.
Exactly where in this range we do notknow.
NUMERICAL MODEL
Rather than using the graph in figure 10-13 to calculate thefraction of the ocean equilibrated for a gi ven mean age for thefossil fuel CO 2 molecules, we will adopt a more rigorous approach.
It invol ves the use of the 1-D numerical model referred
to by workers in the field as the "Oeschger" model (the name ofthe senior member of a group in Bern, SWitzerland, responsible forthe development of this model).
The basics of this model are illustrated in figure 10-14.
The model's ocean consists of 76 boxesof equal volume one above the other.
Each box represents a layer
in the ocean 50 meters thick.
Mixing is accomplished by exchang-

526

~

~

~ 10

0

LC)E 80.........
•

6

a::
a..

4

0

0

....J
<X

::::>

2

Z

Z

«

0

1960

2000

2020

2060

YEAR
60
..........

50

'-

~
""'-"'

W 40
~

« 3-0
z

~ 20
~

10
0
1960

1980

2000

2020

2040

2060

YEAR
Figure 10-12. Projections of fossil fuel C02 production rates overthe next 80 years for a variety of energy scenarios (upper panel)and the corresponding mean ages for fossil fuel C02 molecules
so produced (lower panel).

527

VENTILATiON TIME (yrs)
•

-.J

-:::>

0

w
Z

0

W
J:
f1.1-

0
Z
0
f-

400

900

(J)

0.8

«W
U

100

0

1.0

0.6

a:::
~
w
>0
co

.-x

w
z
a:::

0.4

NATURAL
RADIOCARBONBASEDEQU I LIBRArlON
TIME FOR
DEEP SEA

0

u.

w
z
~
a:::

C>

0.2

U

<t

TRITIUM CALIBRATION POINT0::

u.. 0.0

0

10

20

I

30

{VENTILATION TIME)/2
10-13. Plot of fraction of ocean volume equilibrated
against ventilation time.
One calibration point is based on the
distribution of bomb tritium in 1973 and the other one on thesteady state distribution of natural radiocarbon.
The straightline is drawn assuming that the fraction of volume equilibrated
with fossif fuel C02 increases with the square root of the mean
age of the fossil fuel C02 molecules.
The range limits are forventilation times corresponding to mean C02 ages of 30 and 60
years (the conversion factor from mean C02 age to equivalent ocean
ventilation time is 0.57).

Figu~e528

POLAR00::

l1Jl1J~?:il---+---+'--------Iw~--'I""'~~"""'"::E..J

OUTCROP2

1.~~--I~---~~4-f~~'"
3
4
5

x- I
OTI--+--+----------.....~i-t
00

CX)lO

...

x

rt").i.I-......~--4--------

x+ I- - _ 4 r _.......

73

74
75
76
Figure 10-14. Multibox ocean model used to calculate the uptake offossil fuel C02 by the ocean.
The ocean is divided into 76 layers
each wi th the same volume.
Water is traded between adjacentreservoirs (equivalent to vertical eddy diffusion).
Water also
passes up from the bottom of the model to the level designated as
the source for new deep water. From here it is transferred to thepolar outcrop where it is cooled without undergoing modificationin its Le02 content and carbon isotope ratios.
This water is thentransferred to the model's bottom box.
The PO q reaching theuppermost reservoir is entirely utilized by plants.
These plants
are assumed to incorporate 105 carbon atoms for each phosphorus
atom.
The organic matter so produced falls toward the bottom.
As
it falls it is progressi vely oxidized returning the carbon andphosphorus to solution.

529

ing water between these boxes (see left hand arrows).
Bottomwater formation is accomplished by transferring water from one of
the intermediate depth boxes (consistent with our discussions of
the origin of deep water in chapter 8) to a polar surface ~ox,where it is cooled.
The assumption is made that no EC02 or 1 C/Cchange occurs during this cooling.
This water is then added tothe bottom box from whence it upwells through the series of boxes(right hand arrows) completing the cycle. Organic matter producedin the surface water settles toward the sea floor.
As it falls it
is progressively consumed by animals and bacteria.
Thus thismodel includes all the basic first-order processes going on within
the sea (with the exception of CaC03 formation and dissolution).
Being a one-dimensional model, it averages the lateral inhomogeneities in the sea.*
Our task now is to calibrate the parameters in the model.
They are as follows:
1)
the C02 exchange rate between the atmosphere and surface.
ocean
2)
the exchange rate of water between the ocean reservoirs
3) the flux of deep water
4) the depth in the ocean from which new deep water is drawn
5)
the rate of rain of organic matter from surface water
toward the deep sea
6) the fraction of this debris oxidized at each level in themodel*It must be pointed out that the modifications we have made in
the Oeschger model are more of a cosmetic than a substantive
nature. First, we use a somewhat hi~her vertical mixing rate ~nthe upper part of the ocean (1.6 cm /sec compared to 1.25 cm /sec).
We do so based on the global tritium data set, which was
not available at the time Oeschger and his colleagues wrote
their paper. As the uptake of fossil fuel C02 by the sea varies
with - square root of the diffusivity, this adjustment does notchange things very much. Second, we include the cycle of particulate matter in the ocean. The Oeschger model has received unjust criticism for not including this mode of transport.
The
addition of fossil fuel C02 to the oceans should not significantly alter the amount of carbon transported by the ocean'spart iculate cyc les (as we have seen, they are 1imited by theavailability of P and N).
Thus inclusion of particulate cycles
does not change the amount of foss i 1 fuel CO 2 taken up by themodel ocean.
The third change we have made is the inclusion ofdeep 'water production.
Because we do not· permit EC02 or 14C/ Cchanges in the polar reservoir, this change has little effect on
the 30 to 60 year time scale of importance as over the next 100
years.
The loop we have inserted merely provides an alternatemeans of stirring the deep sea, which yields more reasonable
deep water profiles for 14 C/ C and temperature.
We decrease the2vertical diffusivity from 1.6 cm 2/sec to 0.5 cm /sec below 700
meters depth to compensate for the mixing generated by the deep
water loop. Thus, while our modifications of the Oeschger modelyield more realistic vertical profiles, they do not substantially change the amount of fossil fuel C02 taken up by the ocean.

530

We accomplish this by forcing the model to meet the following
constraints:
1) It must carry tritium to the proper mean depth (i.e., 375m
in 10 years).
2) It mus t produce a na tural 14 C/ C vs. depth relationshipconsistent with that observed in the ocean.
3) It must match the input of natural radiocarbon across the
air-sea interface with radiodecay within the sea.
4) It must produce the appropriate vertical profiles ofphosphate.
The first step in this calibration is to select a set ofdynamic parameters which will simultaneously fit the distributionsof bomb tritium and of natural radiocarbon.
We found by trial anderror that the following values yield satisfactory fits to both ofthese data sets:
Vertical diffusivity
100 to 700 meters

1.6

cm 2
sec

Vertical diffusivity
700 to 3800 meters

0.5

Deep water upwelling rate

4.4 m/yr ( 50 Sverdrups)

Depth of origin of deep water1000 meters

cm 2
sec

In order to get the preanthropogenic 6 14 C for surface wa~er (i.e.,-50%0) the COinvasion rate must be 17.5 moles/myr (forEf atm).
PC0
=
280
x
102
The rate of rain of organic matter from the mixed layer isobtained by multiplying the flux of phosphorus to the mixed layer(in moles/m 2 yr) by 105.
The mixed-layer P0 4 content is maintained at zero.
The model ocean is given a total amount of phosphorus equal to that in real ocean.
Then by trial and error thedepth distribution of the destruction of organic residues isadjusted so as to generate a steady state P0 4 depth distribution
similar to that found in the sea.
It is assumed that no P04 is
lost during the passage of water through the polar reservoir.
The
proper depth of distribution of P0 4 achieved using the relationship:
Rz = R1 exp [ _ _z ]577
where R1 is the organic residue oxidation in the first box
beneath the mixed layer and z is the depth in meters.
R1 isequal to 8.5% of the total settling rate of organic residues fromthe surface reservoir.
As can be seen in figure 10-15, the natural radiocarbon profile gives a reasonable representation of the horizontally averaged ocean distribution for this isotope.
Surface water has a531

P0 4 (fLmol/kg)
0°:-

....

-..2

--.3

::c

2
....
a..

w

Q

AVERAGE P04 = 2. f ,umol/kg
AVERAGE ~14C = -155%0

4 ....- - - - -....- - - - - -.....- - - - -..
-200
-50
-100
-150

il l4 C (°/00)
with depth of phosphate and t:. 14C for
the modified Oeschger model described in the text.
The parameters
used in these calculations are given in the text
These curves
are both reasonable approximations of the actual horizontally
averaged vertical distributions of these properties.

Figure 10-15. Distributions

0532

1.0 r-----.----~--___,~--.,..._--_...--

0.8

E

,
I
\

\

0.6

+00

\
\

,

~

INSTANTANEOUS
CO 2 INPUT

\

E

+-

o

0.4

~

.......

\\

f" . . .

0.2

INSTANTANEOUS
14c INPUT (OCEAN ONLY)INSTANTANEOUS "14c INPUT(OCE AN+ FORESTS +SOILS)

0.0 ~--------_......_ _-.a..

a20-- -- - - -a.-_ _- L ._ _- - - '40

60

TIME (yrs)10-16. The fate of a single sudden input of C02 (upper
solid curve) and of 14 C (lower solid curve) to the atmosphere, aspredicted by our modified Oeschger model.
The decay of these socalled "del.ta" functions is represented by the fraction of theamount of the original amount remaining in the air at any gi ventime after the injection.
The lower dashed curve shows the decrease in 14C if dilution by the soil and forest reservoirs areincluded.
As these reservoirs will not take up any C02, the upper
curve 1s not altered by the addition of these reservoirs in themodel.

Figure533

~14C value of -500/00.
The 14 C/ C value decreases across themain oceanic thermocline reaching a minimum of about -190%0 at
a depth of 2 kilometers.
Below this depth it rises, reaching avalue of -160%0 at the model's bottom.
The mean ~14C for themodel ocean (i.e., -155%0) is close to that for the realocean.
The model yields a mean penetration depth for bomb-produced
tritium of 374 meters as of 1973.
The mean residence time forbomb tritium atoms in the ocean at this time was 9.9 years.
As shown in figure 10-15, the depth profile for P0 4 is areasonable representation of the horizontally averaged distribution of this property in the sea.
Surface water is phosphate
free.
The P0 4 concentration increases steadily through thethermocline, reaching a maximum of 2.8 x 10- 6 moles/kg at a dept~of about 1.3 km. The value then drops, reaching about 2.35 x 10moles/kg at the bottom of the model ocean.
The response of this model to instantaneous injections to theatmosphere of bomb radiocarbon and of fossil fuel C02 is shown infigure 10-16.

CROSSCHECKS ON THE VALIDITY OF THE NUMERICAL MODEL
Fortunately in fixing the parameters of our one-dimensional
ocean model we did not use all the available information.
Thatleft over can be used to check the reliability of the model. Fourtypes of information are available for this purpose.
First, wehave the results of Keeling and his coworkers who have kept trackof the CD2 pressure in the atmosphere from 1958 to present.
Theselong records are available from two places: high on an extinct
volcano on the is land of Hawaii and on the continent of An tarctica.
The Hawaii record is shown in figure 10-17.
The CD2 partial pressure at Mauna Loa shows an increase of 21 x 10- 6 atm between 1958 and 1978.
It is this difference which provides one of
the crosschecks on the model.
The modified "Oeschger lf model presented above yields an increase over this time interval of 24 x
10- 6 atm.
The calculation is carried out by stepping the modelthrough time.
At the start of the calculation the C02 distribution in the system is at steady state.
This steady state is perturbed by the year by year addi tion of excess CO 2 to the model's
atmosphere (following the historic sequence as shown in figure10-1).
We speak of the increase in PC02 over this twenty-year
period rather than the absolute values because of the uncertainty
in the CO 2 .concentration in the atmosphere prior to the industrial
revolution.
Unfortunately, we have no measurements of sufficientaccuracy or reliability prior to 1958 to permit the atmosphericrecord to be extended back in time.
Instead, by adjusting the
pre-industrial atmospheric C02 content, we force the model toyield the atmospheric CO 2 content observed in 1958. This approachis not entirely adequate, because it neglects the contribution of
the forest and soil CO 2 .
We can get an idea of the influence of the C02 from forestsand soils by running the scenario shown in figure 10-2 through themodel.
As we have as yet not shown why this scenario is to be

534

340
336
332
-E 328
0.
0.

-o

C\I

u

324

320
316
312

340
336

-

332

E
a.
a. 328-o 324
C\J

u 320
316

312

Figure 10-17. Plot of atmospheric C02 partial pressure (in dryair) for the period 1958 to 1980 as measured on the island of Hawaii.
The upper panels show weekly means. The lower panel shows
running means designed to remove the annual cycle.
The strong
seasonal cycle in the Hawaii record represents seasonal uptake andrelease of C02 by the terrestrial vegetation in the northernhemisphere.
The station at Mauna Loa was established by Charles
Keeling of Scripps Institution of Oceanography (556,558,562a).

535

preferred, the reader will have to bear with us until this is donein a few more pages.
As can be seen in figure 10-18, the forestplu~ soil CD2 input raises the atmospheric CD 2 pressure by 49 x
10- atmospheres as of 1980.
However, because of the decline in
the input of forest and soil CD 2 over the last 50 years (see figure 10-2), the change in atmospheric CD2 content produced over theperiod 1958 to 1978 by the forest and soil CD 2 contribution alone
is only about 2 x 10- 6 atm.
Hence, while the decline in forest
cover and in soil organics has influenced the overall buildup of
atmospheric CO 2, it has not played a major role over the 1958 topresent period.
When the 2 x 10- 6 atm rise in PCO from forests and soils
is added to the 24 x 10- 6 atm rise from tossil fuels, a total 0t
26 x 10- 6 atm is obtained.
This is 20% larger than the 21 x 10atm increase observed at Mauna Loa during the same interval oftime.
Thus our model does not take up enough CO 2 to pass thiscalibration test.
Before discussing why this might be the case,let us consider the other crosschecks.
The second potential crosscheck comes from the 13 C/12 C record for atmospheric C02.
As both fossil fuel and biomass carbon
have o13 C values averaging -260/00 (compared to that of about-6 0 /00 for the preindustrial atmosphere), the addition of excess
C02 to the atmosphere from these sources should have driven down
the 13 C/12C ratio in atmospheric C02.
As is the case for the qq2
content of the atmosphere, the magni tude of the atmosphericCchange will depend on the amount of C02 released and on the extent
to which this C02 has equilibrated wi th ocean carbon.
ThEF3e ~~,however, a difference in the C02 pressure response and theC/ Cratio response.
Any atmospheric carbon isotope anomaly will
eventually be diluted in accord with the ratio of ocean carbon toa tmosphere carbon (i. e., about 60).
As we have seen the corresponding reduction for fossil fuel CO 2 wi th the ocean after fullequilibration will only be one-tenth as large (i. e., about 6).
F.urther, the carbon in living biomass and in soils will also dilute the carbon isotope anomalies.
On the other hand, these reservoirs will not take up a significant amount of the excess CD2.
The living biosphere and soil reservoirs are incorporated
into the model as follows.
We take the mass of living biospherecarbon to be equal to that in the preindustrial atmosphere (i.e.,
50 x 10 15 moles) and that of soil carbon to be three times that in
the preindustrial atmosphere (i.e., 150 x 10 15 moles).
We assumethat 30% of the living biosphere consists of leaves, rootlets, andsmall plants which are replaced annually.
The remaining living
biosphere is assumed to be wood.
Half is gi ven in replacement
time of 20 years and the other half a replacement time of 100
years.
We assume that soil carbon is replaced on a 250-year timescale.
The difference in isotope effects between the models run
with and without the biospheric reservoirs is about 20% (seefigure 10-16).
We can get an idea of the expected size of the 13C anomaly
as follows.
As of 1978 the amount of fossil fuel carbon released
to the atmosphere was equal to 25% of the amount of CO 2 in thepreindustrial atmosphere.
A mixture of 100 units of carbon with536

r-----r---.---..........---....--....--......- -..... 350
100....... 80.......

E

E

+-

0

+0-

eD

b

0
(0

60

J

0

'-""

'-"

N

0

u

N

0

40

u

o..

C-

<J
20

250
0
1800

1850

1900

1950

YEAR
Figure 10-18. Atmospheric C02 partial pressure reconstruction madewith the modified Oeschger ocean model described in the text.
Both fossil fuel C02 (as estimated from historic records) andforest - soil C02 (as given in the scenario of figure 10-2) were
input to the model.
The contributions of the two C02 sources areshown separ.ately.
The curve forthe total is constrained to go
through the value of 315 x 10- 6 atmospheres obtained by Keelingand his coworkers for the year 1958.

537

a o13 C of -60/00 (the pre~~thropogenic value for atmosphericCO 2) and 25 uni ts wi th a 0 C of -26 % 0 (the value for mean
fossil fuel and terrestrial biosphere CO 2 ) would have a o13 C of
-10%0.
Thus in the absence of any dilution with carbon ff~mother reservoirs the atmosphere would have decreased in its 0 C
value by 4%0 from the fossil fuel C02 contribution alone.
The
model is run for 13 C the same way as it is for CO 2 • Fossil fuelC02 and forest and soil C02 are added to the model atmosphere in aseries of time steps.
Mixing of carbon between the atmosphere andocean occurs diluting the atmospheric 13C anomaly generated bythese anthropogenic additions.
Our model, when run for fossil fuel CO 2 alone, predicts ao13 C drop of 0.580/00.
When the forest-soil scenario of figure10-2 is included, this result decreases to 0.51%0.
In thiscase the continuing dilution of the earlier large forest and soil
input more than compensates for the curre~t 1~lease of CD2 fromthese sources.
The earliest accurate 1 C/ C ratios made on
atmospheric CO 2 samples were carried out during the mid-1950's.
Keeling, who conducted these early measurements, provided samples
collected during 1978 to Mook of the Groningen Laboratory in the
Netherlands for c?mparison analyses.
They concluded from thiseffor t th a t th e 0 3C dr op p e d by 0 . 55 00 be t wee n 1956 and 19 78 .
As the uncertainty in the atmospheric change over the 1956 to 1976
period is probably of the order of 0.15 0 /00, the agreement
between observation and prediction is adequate.
It is now time to show how we obtained the forest-soil
scenario in figure 10-2.
It was deri ved from a consideration ofan extended atmospheric 13C record obtained by measurements on thecellulose from tree rings.
The idea behind this approach is thateven though there is a large fractionation between cellulose(-26%0) and atmospheric C02 (-6%0), the mean o13 C in cellulose must change with changing atmospheric o13 C . The question iswhether this message can be separated from other sources of variation.*
Freyer at the KFA in Julich, West Germany has conducted extensive research on the 13C/12C record kept in tree rings.
Afterfinding that trees growing within forests did not give the expected trend, Freyer concentrated his efforts on solitary trees.
Heobtained such trees from the French Atlantic coast, from the Spanish Pyrenees, from North C~rolina in the U.S., from the Eifel andBlack Forest in West Germany and from northern Sweden and Norway.
The resul ts on the Swedish Scots pines, which go back to 1530,
suggest that the atmospheric 13 C/12 C ratio was quite constant°/*Climatic changes will alter conditions at the leaf surface andhence change the magnitude of the isotope fractionation which
occurs during photosynthesis.
Canopy effects are produced by
the passage through the canopy of the C02 produced by the decayof vegetf~ion in the soils beneath the trees.
Reincor~orationof thisC depleted C02 would further reduce the 13 C/l C ratio
in the forming cellulose.
Thus changes in the geometry of the
canopy due to growth and death of neighboring trees and due tothe growth of the tree of interest could lead to changes in theisotopic composition of the cellulose formed in the tree.

538

--..

--

- -

0

J-~

0

0~

-u

rt)

. • -_.-.".
-.... _~~-- - ~

--

-

-I

GO

<l
-2
1500

1600

1700

1800

1900

2000

YEAR

•
o a o~··.

00.

a 0 ~<»

c!' 0 0o.

-I•a• •eeee

-2
1800

1850

1900

1950

YEAR
Figure 10-19. Average trend for the last 500 years in atmospheric
1 3 C/ 12 C ra tios as es tima ted from measurements on soli tary Scots
pine trees from northern Sweden and Norway (upper panel).
Comparison of the Scots pine record (open circles) wi th the average
record established from measurements on trees from the Pyrenees,
the west coast of France, North Carolina, and the Eifel and Black
Forests in Germany (closed circles). These measurements were madeby Preyer and his colleagues at KFA in JUlich, West Germany (574).

539

prior to 1830 (see figure 10-19). 'Between 1830 and 1970, a dropof 1.8 0 /00 occurred.
Also shown in this figure is a comparison
for the last 200 years of the Scots pine record with the mean
trend from the other areas mentioned above.
As can be seen, theagreement is quite good.
Before interpreting this record, it must be pointed out that
other workers have not been so successful in finding trees thatgive smooth records.
Stuiver of the University of Washington hasstudied soli tary trees from si tes along the west coast of theAmericas and Francey of Australia has studied trees from Tasmania.
Nei ther has found the trend obtained by Freyer.
Their records show considerable variability on time scales of 100 years andless.
Freyer himself was of course not able to find the expectedrecords in Spessart oak trees from forests in Germany.
Until thereasons for these differences between species of trees and growthsites is understood, the assumption that Freyer's records represent the decline in atmospheric o13 C rests on shaky ground.
Nevertheless, we will proceed on the assumption that Freyer's
curve is a valid representation of the atmospheric record.
Shown in figure 10-20 are the model resul ts* for the inputof fossil fuel C02 alone, for the input of forest-soil CD2 alone
and for the sum of these two inputs.
As can be seen by comparisons with Freyer' s results (black circles in figure 10-20), thefossil fuel effect alone yields only half the amplitude and diff ers considerably in shape from the tree-ring-based atmosphericrecord.
If the difference between the tree-ring-based curve andthe fossil fuel C02 curve yielded by the model is ascribed to theaddition of the forest and soil C02 in the atmosphere, then an
input function for this CD 2 can be constructed which, whencombined wi th the fos s i1 fuel CO 2 input curve, will yield thetree-ring-based curve.
This is the origin of the scenario in
figure 10-2.
As shown in figurEf 10-20, the combined forest-soiland fossil-fuel input yields a 0 3 C record in accord with Freyer's
13 C: record.
A possible crosscheck on this forest + soil scenario would beto reconstruct the o13 C: record in the surface ocean using measurements on ring-dated corals.
The trends expected from fossil fuelC02 alone and from the combination of fossil fuel C02 and forestsoil C02 is shown in figure 10-21. The single record available todate (i.e., that on a Bermuda coral) is shown for comparison.
When faced with our forest scenario, Mook of the Netherlands
made an interesting observation.
He pointed out that the o13 C
values for. the surface ocean as of 1973 (see figure 6-12) showed
the ocean as a whole to be nearly at steady state with the atmosphere.
He asked if our combined forest-soil and fossil fuelscenario would disrupt this steady state (i.e., would the surface
ocean and the atmosphere have been at stable carbon isotope steadystate prior to 1850).
As shown diagrammatically in figure 10-22,
if anything, the agreement would have been better at this time.
*In these calculations we use the kinetic fractionation factor
for C02 invading the sea of 20 /00 suggested by Siegenthaler
and Munnich rather than that of 14%0 observed for C02 uptake
into strong alkali solutions.

540

o....--...

--~:::--_FOREST + SOIL........
0~0.........

u

-I

•

rt')

t:()

<l

••

-2

1800

1900

1850

1950

2000

YEAR
Figure 10-20. Atmospheric o13 C response calculated using themodified Oeschger model described in the text.
As for the C02
partial pressure reconstruction, both the contributions of fossil
fuel C02 and forest-soil C02 are included in the calculation.
The
contributions of the two sources are given separately. As was thecase for C02 partial pressure the slope over the last 20 years is
not strongly influenced by forest-soil inputs.
The trend established through Freyer's
tree-ring measurements is shown forcomparison.

541

°r---~~::---;;r----...

o

-

~

U
rr>

GO -I

SURFACE OCEAN<J

13 C TREND
1800

1900

1850

1950

YEAR
0 ....- -....- -....- -.....- - - - - - -.....- -.....- -....

-

o
~
o

u

rt')

(,() -I

BERMUDA CORAL

<l
1800

1850

1900

1950

YEAR
change in the 0 1 3 C for average surface ocean
water as calculated using the modified Oeschger model described in
the text.
The contributions of fossil fuel C02 and forest-soil
releases are shown separately.
In the lower panel the resul ts
obtained by Nozaki and his associates at Yale Uni versi ty on a
Bermuda coral (14) are shown for comparison.
Figure10-21. The

542

........
0

~
0

3

........
Z

<t
w
u
0

w
u

2

~
0::

:::)

en
u

-

rot')

c.<)

o -----.&.-----.....-----..
10
20
30

0 .....

TEMPERATURE (OC)
Figure lO-2~~ Diagrammatic representation of the relationship between the 0 C for surface water carbon and temperature as measured in 1973 (see figure 6-12 for data) and the 013 C which surface
water carbon r~uld have had were it at equilibrium with the 1973
ar~osphere '(0C = -7.2%0).
Also shown are the surface watero C averages reconstructed for the period before the impacts ofman (the combined forest-soil and the fossil fuel C02 input
yielded a drop of 0.95%0 in the model's surface ocean between
1850 and 1973).
Finally, the 013 C values for ocean water at
equilibrium with the pre-18
atmosphere (013 C = -5.6%) is shown.
The 1.6 0 /00 decrease in 013 C for atmospheric CO 2 between 1850
and 1973 is that obtained from the model. While individual areasof the ocean were not at isotope equilibrium with the air the sea
surface as a whole must have been at steady state prior to theyear 1800.

°

543

Our model suggests that the o13 C for average surface ocean water
was 0.1950/00 higher prior to 1850 than it was in 1973 and thatthe 0 3C for the atmosphere was 1.6 0 /00 higher.
Making these
corrections, the crossover between the equilibrium curve comes at
20°C as opposed to 14°C in 1973.
As the mean surface ocean temperature is 18°C, the pre-1850 si tuation is, if anything, an improvement.
While this cannot be taken as support for the forest
hypothesis, at leas t it does not lead to difficul ties in thisregard.
The third crosscheck comes from the decrease in the atmospheric 14C/ C ratio between 1860 and 1950.
As fossil fuels carryno 14C , input of CO 2 from this source lowers the atmospheric ratio.
On the other hand, forest and soil carbon have 14 C/ C ratios
reasonably close to the atmospheric value, hence input of C02 fromthis source would have no significant effect on the atmospheric
14C/C ratio.
We must terminate this comparison in 1950, becausesoon thereafter the input of bomb 14C to the atmosphere became
significant.
Were no mixing of atmospheric, ocean and biospherecarbon occurring, then a decrease in 14C/12C ratio of 10% would be
expected.
The atmospheric 14C/ C record is preserved in treerings, and in this case the small changes in the fractionation
factor, which bedevil the 13C reconstructions, have no significant
effect. The measurements of Stuiver of the University of Washington on tree rings from 1860 to 1950 are shown in figure 10-23.
A
2% decline is found between 1850 and 1950.
Also shown in thisfigure are the trends predicted by our modified Oeschger model.
The agreement is excellent.
In comparing the observations with the model curves, it must
be kept in mind that during the pe riod of the 14C/ C decline,chan~es in the production of 14C similar to those which affected
the4C/12C ratio of the atmosphere prior to 1850 (see figure 732) must also have been operative.
Minze Stuiver, working withhis associ~te Paul Quay, has developed a relationship between sunspot ac ti vi ty and 14 C production which accounts for much of thevariability in the 14C/ 2C ratio for atmospheric C02 for preindustrial times (19).
Based on the record of solar acti vi ty between
1850 and the present these investigators have reconstructed thecosmic ray production of 14 C over the past 100 years.
They conclude these changes have led to a decline in the natural 14C/ C ratio of about 3.0 0 /00 during the last century.
They recommendthat this amount be deducted from the observed change when looking
for the change due to fossil fuel CO 2.
Thus the fossil fuel CO 2
induced decrease between 1850 and 1950 appears to have been about
18 0 /00.
As can be seen from figure 10-23, the model predicts aslightly larger change.
However, because of the uncertainty inthe correction for the variabili ty in the cosmic ray production
rate, the agreement between the observed and model curves isadequate.
The surface ocean should also show a Suess effect due to theaddition of CO 2 free of 14C • As shown in figure 10-24, the model
predicts an 8 0 /00 between 1850 and 1950.
The reader is referred
to figure 7-34 which shows the results on corals. While an effectof about the right magnitude is seen, the experimental errors on
the measurements are too large to yield a meaningful tes t of the

544

-

o

o~

-u

-10

v

<J

•

'--_.L..._,S.LS-O-.L-_...a-_......._ .......-1-9...O-O-....._ ......_..a.-_-..--19fO

YEAR
Figure 10-23- CDmparison of the model derived ~14C decrease causedby the addi tion of 14C-free fossil fuel CO 2 to the atmosphere(i. e., the Suess effect) wi th the measurements of Stui ver of the
University of Washington on tree rings covering the period 1800 to1950 (26). .

545

,...".

0

a

0~

""'"

u

~

-<l
<J

SURFACE OCEANSUESS EFFECT
10

1800

1850

YEAR

1900

1950

Figure 10-24. Model calculation of the ~14C reduction expected forsurface ocean carbon due to the addition of 14C-free fossil fuelC02 to the atmosphere.

546

model.
The
fourth
and
last
of
the
crosschecks
comes
from
measurements of the distribution of the 14 C produced during bomb
testing.
In particular we are interested in comparing the
distribution of the 14C between the atmosphere and the sea
predicted by the models wi th that observed at the time of the
GEOSECS program.
The amount of bomb- 14 C in the ocean at the timeof the survey can be estimated from the water column inventory
data discussed in chapter 8.
As summarized in table 10-5, 314 ±35 x 10 26 atoms were in the sea as of 1973.
Our model is then run using the observed atmospheric 11 14 C
curve from 1950 to 1973 (see figure 8-3) to predict the number ofbomb 14C atoms taken up by the ocean.
In this case the atmospheric 14C/ C ratio measured in atmospheric CO 2 samples taken over
the time interval 1955 to present is used for the calculation
rather than the rate at which bomb 14C atoms were added to the atmosphere.
For 1973 the model gives 302 x 10 26 bomb 14 C: atoms in
the ocean.
The agreement between the predicted and observed values is excellent!
This same ocean uptake could be achieved with several different combinations of gas exchange and vertical mixing rate.
To
see if we have the right combination another check can be performed.
The average excess 11 14 C for surface water in 1973 can be compared to the model calculated value.
If, for example, our gas exchange rate were too high and our vertical mixing rate too low, we
could get the right bomb 14C inventory but too high a surface water excess. The 11 14 C increase (over natural levels) for surface
water as measured during the GEOSECS program is summarized in table 10-6. The average is 160 ± 15 0 /00.
Our model yields a prediction of 172 0 /00
in satisfactory agreement with observed
value.
In summary, then our model satisfactorily predicts the carbon isotope anomalies due to fossil fuel C02 release and due tothe invasion of bomb _14C into the ocean.
On the other hand, it
seems to underpredict the uptake of fossil fuel C02 by theoceans.
This apparent disagreement between isotopic and chemical
response may stem from the one-dimensional character of themodel.
To understand why this is the case we must turn our
thoughts back to the outcropping isopycnals of chapter 8.
An alternate model could be made which would permit entry to the main
thermocline directly from these outcrops.
This model could thenbe calibrated so that it gave the right distribution of naturalradiocarbon, the right Suess effect and the right bomb- 14 C inventory and bomb- 14 C excess in surface water.
When such a model isrun for fossil fuel C02 uptake, it will likely yield a greater
ocean uptake than would be obtained by the model used here.
Thereason is tqat the resistance to entry through the outcrops isgreater for 4 C than for CO 2 (remember we showed in chapter 3 thatthe equilibration time for C02 was an order of magnitude shorter
than that for carbon isotopes).
Siegenthaler of Bern (699) has
done some preliminary calculations designed to get some idea ofthe magni tude of this enhancement.
Their resu1 ts show that asmuch as 20% more C02 would be transferred into the sea.
If so,
the discrepancy noted above for C02 uptake would be largely

547

Table 10-5. Summary of the bomb- 14 C inventories for various zones of theocean (as of 1973). Based on measurements made by Stuiver of the Universityof Washington and Ostlund of the University of Miami (21,22).
Atl.
Antarctic (>45°8)

~5

Pac.
Mean Global
9
10 atoms/cm 2

Frac. Ocean
Area

Excess 14 C
10 26 atoms

'"'9

'"'7

0.167

~40

s. rremp. (45°S-15°8)

9.5

9.7

9.6

0.264

92

Equator (15°S-15°N)

3·3

5.4

5.0

0.281

51

N. Temp.

(15°N-45°N)

15.3

14.2

14.6

0.192

101

N. Pac.

(>45°N)

'"'6

-

0.028

'"' 6

N. Atl.

(>45°N)

-

0.036

19

0.032

'"' 5

CJ1

~

00

Arctic Ocean

18.6

-

-

-

'"'4

314±35

Table 10-6. Summary of surface bomb- 14 C excesses for surface wa-

ters (as of 1973). Based on measurements made by Stuiver at the
University of Washington and Ostlund of the University of Miami aspart of the GEOSECB program (21,22).

RegionNo. ofStns.

GEOSECS
14
6
C

Prenuclear
14
6
C

Excess
14
A C

0/00

00/0

00/ 0

ATLANTIC

N. Polar

8

50

-70

120

N. Temp.

8

145

-50

195

Equa.

6

85

-65

150

S . Temp.

6

126

-50

176

Antarctic

6--. . 100

PACIFIC

N. Polar4--. . 130

N. Temp.

15

163

-50

213

Equa.

5

61

-70

131

S . Temp.

7

141

-50

191

Antarctic

4--....115WHOLE OCEAN
N. Polar(0.096)*

125

N. Temp.

(0.192)*

205

Equa.

(0.281)*

140

S . Temp.

(0.264)*185

Antarctic(0.169)*110160±15

Mean*Fraction of ocean area549

removed.
It will be a long time, however, before we know enoughabout the ocean to permit reliable 2-D modeling.

PREDICTION OF FUTURE C02 LEVELS
Keeping in mind that our I-D model may slightly underestimate
the uptake by the sea of fossil fuel C02, let us use it to predict
future atmospheric CO 2 pressures for a variety of scenarios.
Ourfirst effort in this regard will be to look at the next 80 years(i.e., 1980 to 2060).
As shown in figure 10-12, in so doing weare not making long extrapolations from the time scales for the
tracers used to calibrate and test the model.
Even if on the average we produce over the next 80 years the same annual amount offossil fuel CO 2 as we did in 1980, the mean age of the C02 molecules will not exceed 60 years.
On this time scale they should
penetrate only about twice as far into the sea as had tritium at
the time of the Atlantic and Pacific GEOSECS surveys.
In making these calculations we raise the rate of C02
invasion into the ocean in accord with the C02 pressure in theatmosphere, i.e.,
I = 17.5

PC0 2 6 moles/m 2 year
280 x 10-We also change the Revelle factor as the C03= ion content of themodel's surface reservoirs falls.
Four cases have been considered.
Man's production of C02 is
assumed to increase at rates of 3%/ y r, 2%/y r, l%/yr, and O%/yr.
The calculation is begun in 1980.
At this time the model has thecharacteristics achieved at the end of our combined fossil fuelforest C02 scenario.
The results are portrayed on a plot of mean
age of C02 versus total amount of C02 produced (see figure
10-25) .
As can be seen, the predicted PC0 2 depends mainly on
the amount of C02 produced.
The question comes to mind why the atmospheric C02 content is
so insensitive to the time history of the input.
To see why thisis the case let us compare the two points on figure 10-25. One isfor the no-growth scenario in the year 2060.
The other is for the
3%-growth scenario within the year 2020.
For both points the C02
input is abou t 4.8 x 10 16 moles.
The atmospheric PC:o 2 for theno-growth point is 450 x 10- 6 atmospheres.
For the 3%/yr growth
it is 475 x 10- 6 atm.
The excess PC02 over that in 1980 (340 x
10- 6 ) is 110 X 10- 6 atm for the O%/yr point and 135 x 10- 6 atm forthe 3%/yr point.
The difference in mean CO 2 age (and hence inocean residence time) for the two cases is a factor of two.
Hence
the penetration depth difference should be the square root of
two.
As in both situations roughly twice as much C02 remains in
the air as it enters the sea, an increase of 1.41 in the ocean uptake will lead to only an 0.80 decrease in the extent of atmospheric buildup.
This agrees wi th the actual reduction gi ven by
the model (i.e., 110/135).
Thus we see that at least within thecontext of our model extrapolation to the future can be done quite
reliably using figure 10-25.
Any scenario (exponential or otherwise) yielding a mean residence time in the range 30 to 60 years,

550

~

~O%/yr

r..

,.--.
~

~

(\J

~

0

U

I

o

--I
W

<:PO

.

I ...
I ..

'-'

i( IO/o/yr

-

~o

(J)
Cf)

...
°

/ .....

..

I

"'..

..· /
'
o
·
·

0

<t:

Z
<t
W
~

40
_"'Ii

'-0

I ... /
I 1.
0

0

0 ".

o

...

..

.....

'. ....

.....

(0...-:,.... .,.."

...

oI 0

0

.....

IL.

W
C>

...

o~

I ...

--I

...
..

". /

f..

LL

/

.~

1tQ1

:::>

~

~. •••
0

..

g

...

0··.

/ .

t

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0
0

Q)

(l)

f'...

__ 1../
L__~·~·l~~·/yr~~ /1~ - - - -"oe.
.
:.
30
V

t

o -

2

/

\.

4

••••• ••••

6

8

10

12

14

AMOUNT OF FOSSIL FUEL CO 2 PRODUCED (IO

16
I6

18

mol)

Figure 10-25 .. Contours of C02 partial pressure predicted for the
future as a function of the mean age of fossil fuel C02 moleculesand of the total amount of fossil fuel C02 generated.
These contours were computed by inputting to the model C02 scenarios
involving 0, 1, 2, and 3%/year growth rates for fossil fuel C02
emissions.'
The traverses for each scenario are shown by thedashed lines. Time lines (dotted) are shown.
While the scenarios
used to cons truct the diagram are based on exponential growth,these diagrams are also applicable to more complicated scenarios
yielding mean age - total production combinations falling wi thin
the field of the diagram.

551

the PC02 read from this graph should agree wi thin 2% of that
calculated by running this particular scenario through the model.
We must hasten to add that this should not be taken to meanthat the predictions given here are accurate to 2%.
As mentioned
above, the model itself is not entirely adequate.
While matching
the isotope effects, the model ocean appears to take up less C02
than it should.
A model which satisified the isotope constraints
and fit the C02 rise over the period 1958 to 1978 would predict
CO 2 increases in the atmosphere over the 1980 level somewhat lower
than those shown in figure 10-25.
As an example of what might happen on a longer time scale wehave run the fossil fuel use scenario shown in figure 10-26.
In
this scenario we assume that fossil fuel use will continue to rise
until the latter part of the next century at which point it willreach a value of about 2.5 times its present val ue.
After thatthe usage is assumed to fall returning to the present level in
about the year 2220.
As of the year 2300 about two thirds of thepresent reserves are assumed to have been consumed.
While quite
arbitrary in its construction this scenario is not unreasonable.
The atmospheric C02 pressures predicted using this scenario(see figure 10-26) continue to rise for more than a century after
the peak in fossil fuel CO 2 usage and then levels off at just over
900 x 10- 6 atmospheres.
Shown in table 10-7 are the carbonate ion contents of surface
water and the isotopic composi tions of atmospheric C02 for this
scenario.
The carbonate ion content of surface water decreases by
a factor of two between 1980 and 2300.
The <S 13C value reaches aminimum in about 2175 and then slowly rises.
Finally the 11 14C
value drops below its prenuclear value reaching a minimum in theyear 2200.
After that it begins to rise.
SOLUTION OF SEA FLOOR CALCITE
A calculation similar to the simplified one we did for theatmosphere-sea equilibrium can be carried out for equilibrium between air, sea and the calci te in marine sediments.
The equilibrium equation for reaction 10-2 is as follows:
a.K =

10-30
PC02

Since the amount of Ca++ in the ocean is about 5 times higher than
the amount of HC03-, as an approximation, Ca++ can be assumed toremain constant.
Hence:[ HC 03-J ~[ HC0 3-Jo

10-31

As most of the carbon dissol ved in the sea is in the form of bicarbonate ion we get:

552

10

W

t-

TOTAL PRODUCTION
AS OF 2300
21.5 x IOl6 mOI

ed:

a::

Z~

o

8

~

"0

t-

6

E

U

::::>q-Q-o0(}:.::;
0N

o

4
2

U

.....__....__....__....__.a.-_ _.a..-_.......

O~-_

1950

2050

2150

2250

YEAR
1000 ~-"""--""'--"---"'--""'---I"'--""........

E

800

+-

<D

0

'0
........

600

C\I

0

u
c..

400

200 .....- .......- -.....- -.....-.-...-----------......- .......
1950
2050
2 ISO
2250

YEAR
Figure 10-26. Possible fossil fuel use scenario to the year 2300
(upper panel) and the atmospheric C02 partial pressures calculated
for this scenario using the model discussed in the text.

553

Table 10-7. Changes in the carbonate ion content of surface ocean
water and in the o13 C and 6 14 C values for atmospheric C02 associated with the fossil fuel use scenario shown in figure 10-26.

Year

Total
FossilFuel CO 2
Production

PC0 2
Atm

Surf
Water
[ C03=J

o13 C
Atm
CO2

14 C
Atm
CO2

10 16 moles

10- 6 atm

10- 6 moles

%0

%0

6

kg

1800

0.0

249

253

-5.6

0

1950

0.5

308

226

-6.8

-24

19~0

1.4

340

212

-7.3

285

2000

2.3

369

202

-7.6

95

2050

5.7

470

172

-8.6

-36

2100

10.7

632

138

-9.6

-Ill

2150

15.3

778

118

-10.0

-141

2200

18.6

864

108

-10.0

-147

2250

.20.5

goo

104

-9.7

-143

2300

21.5

g04

104

-9.5

-135

554

10-32The Keeling fraction can be written as :[ PC0 2 -°pO CO 2]p CO 2

fK[PC0 2 -opoC0 2

p CO 2

]

M°atm

°

M atm

+t[

10-33
LC0 2] [LC0 2 Jo

[l:C0 2 ]O]°M oceanSubstituting the approximate relationship between PC0
andin equation 10-32, we obtain the following relationship:2

LC02

1

280
1 +

5

[j ::~~2

10-34[ P~02

P CO 2As shown in figure 10-27 the Keeling f raction calculated in thi s
As long as
way does not vary too much with atmospheric PC02.
there is calci te present in the sediment to affect the equilibrium, it remains close to 0.04.
The mass of fossil fuel C02 corresponding to any atmospheric CO 2 concentration can be calculated
from equation (10-14).
This relationship is also shown in figure
10-27.
As_ can be seen, much larger C02 releases are required to
produce a gi ven atmospheric PCOz value if the atmosphere remains
at equilibrium with the calcite In marine sediments.
While the interaction with sea floor calcite can greatly
reduce the amount of fossil fuel C02 in the atmosphere, thisinteraction will take several thousand years.
Let us look into
the reasons for this rather long time scale.
The first point to be made is that the calcite solution process can be effective on no shorter a time scale than that for
deep sea ventilation.
As the calcite to be dissolved lies on the
deep sea f.loor, the fossil fuel C02 laden water must reach the
deep sea before this process can begin.
While in the deep sea
this water is not in communication with the atmosphere; hence itis only when it once again returns to the surface that an impact
on the ocean's fossil fuel CO 2 uptake occurs.
The amount of fossil fuel C02 this water picks up during its second pass against
the atmosphere will depend on the extent to which its alkalinity
was increased through calcite solution during its traverse throughthe deep sea.
The more calcite it dissolves in the deep sea themore CO 2 it will be able to absorb during its contact with the atmosphere.
Thus the minimum time scale on which calci te solution

555

z

.06

I

0
I-

-

u

<C

I

I

I

I

I

-

--

.04

ct:
LL

-

(!)

-.z..J .02 -

-

LLJ
L&J
~

.00

I

300

I

I

I

I

500
700
otm -6
PC02 (10 atm)

I

900

40

200

E 30

150

o ...
0

~

"

't't-

=
E
-0
0

(J)

20

JOO

-

~

~

't-

~

10

50

a

900

Figure 10-27. Plots of the Keeling fraction and mass of fossil
fuel CD2 added to the atmosphere versus the C02 partial pressurein the atmosphere, assuming that the deep waters of the ocean are
maintained at saturation with regard to calcite.

556

could impact fossil fuel CO 2 uptake would be the time scale fordeep ocean ventilation (i.e., - 500 years).
We say minimum because complete neutralization of the fossil
fuel CO 2 contained by a parcel of deep wa ter is not likely to beaccomplished during a single pass through the deep sea.
If, forexample, only 10% of this neutralization occurred during each passthen ten passes would be needed!
The solution time scale would
then be 5000 years!
Clearly then, we must get some idea of theextent of neutralization to be expected for a single pass throughthe deep sea.
It is instructive in this regard to calculate thechange in the C03= content of bottom water necessary to produce
a calcite solution rate one-tenth today's rate of fossil fuel C02
production rate.
About 20% of the sea floor is blanketed withsediment rich in calci te (i. e., o. 7 x 10 18 cm 2) .
Today' s fossil
fuel CO 2 production rate is about 5 x 10 14 moles/yr.
Thus, if excess calcite solution is to compensate for one tenth today's rateof fossil fuel C02 production then calcite will have to dissolve
at a rate of 7 x 10- 5 moles/cm 2yr (i.e., 7 x 10- 3 g/cm 2 yr).
This
is about 7 times today's mean rain rate of calcite toward the sea
floor!
In chapter 2 we showed that C03= ion contents 10 lJm/kg below the saturation value were adequate to dissolve most of theraining calcite.
We can use this result to "calibrate" the rateof the solution process.
The relationship between solution rate,F, and C03= ion content can be written as follows:
10-35
where f is the fraction calcite in the sediment, and a is the constant to be evaluated.
At points on the sea floor where carbonate ion content ofbottom water reaches 10 lJm/kg below the saturation value the calcite content of the sediment drops to about 10%.
To achieve sucha low calcite content requires 'that over 95% of the raining calcite be dissolved.
The solution rate is hence comparable to therain rate (i.e., about 1 g/cm 2 10 3 yrs or 1 x 10- 5 moles/cm 2 yr).
From these numbers we can get the value of the constant a:a ~2
1 x 10- 5 moles/cm yr",~3 x 10- 6

10 lJm/kg

mOl eSjlJ m
2
cm yr

kg

Now let us apply this relationship to the excess dissolution
rate of calcite from sediment by fossil fuel C02.
We can rewrite
the equation 10-35 as follows:

1

F

a

If

10-36

As most of the fossil fuel C02 induced calcite solution will takeplace in sediments which currently lie above the lysocline we ca~take f to be about 0.85Thus for a solution rate of 7 x 10moles/cm 2 yr the undersaturation value becomes:

557

1
('oJ

x

25 lJITl/kg/0.85
Thus for a solution rate of 7 x 10- 5 moles/cm 2yr a reduction in
C03= ion concentration of about 25 ~/kg would have to occur.
Because the sediments to be attacked now lie in supersaturated water, to this would have to be added the current extent of supersaturation (i.e., a to 20 ~m/kg). This calculation tells us thatsizable reduction in C03= ion content is necessary before a dissolution rate comparable to even 10% of the current rates of C02
production is achieved.
--Next let us consider how the reduction in C0 3 = ion content
of the deep sea is related to the amount of fossil fuel C02 in theatmosphere.
In our considerations so far we have spoken only of
the changes in the C03= concentration in surface water.
We now
must consider the factors relating the C03= ion content of deep
water to that for the appropriate source water (i.e., the modifications through respiration and CaC03 solution).
Listed in table
10-8 are the C03= ion concentrations, alkalinities, EC02 contents and apparent dissolved oxygen utilizations of the two major
deep water sources and for waters at a depth of about 3500 metersin the major ocean basins.
As can be seen, although the deep waters of the ocean have been subject to significant amounts of respira tion and CaCO 3 solution, the two processes nearly balance,yielding only a small reduction in the carbonate ion concentration.
To calculate the fossil fuel CO 2 induced drop in the carbonate ion content of the deep water sources, we will assume thatthese waters retain the same alkalini ty as measured during the
GEOSECS program and that they achieve equilibrium with the atmosphere prior to their descent.
Plots of C03= for northern component source water against PC02 in the atmosphere are given in
f igure lO-2~.
We can obtain an upper limi t on the reduction in
C03= ion content for deep water by assuming that these reductions are the same as those for these source waters.*
If, ·as we expec t
the CO 2 partial pressure in the atmospherestays below ~OO x 10-~ atm, then the carbonate ion content of deepsea water should drop by no more than 60 ~/kg.
Such a drop would
induce an excess solution rate of calcite of only a few tens ofpercent of the current rate of fossil fuel C02 production.
This
being the ,case it is safe to say that during each pass through the
deep sea only a relatively small fraction of the excess C02
carried by the deep water will be neutralized by calcite in the*Fossil fuel C02-induced-solution of calcite will raise thecarbonate ion content of the deep waters. Failure of the source
waters to achieve equilibrirnn with the C02 in the atmosphere
before descent will also raise the C03= ion concentration in
the source waters.
Finally, as long as the C02 pressure in theatmosphere continues to rise, the reduction in the C03= ion
content of deep water will lag behind that in the source
waters.
Hence our method of calculation yields a maximum
lowering in the carbonate ion content for deep water.

558

Table 10-8. Carbonate

ion concentrations for deep waters in theocean. Based on a compilation of GEOSECS results by Takahashi andhis coworkers (406).
Water
Type

Alk
).J.eq/kg

LC02)lID/kg

AOD)lm/kg

C0 3=)lm/kgSource waters (after descent)
NCW
SCW

2168
2250

2290
2360

65
105

98
79

Basin waters (at 3500 meters)
N. W.
S.E.
Equ.
N.W.

Atl.
Atl.
Ind.
Pac.

2340
2363
2417
2420

2195
2224
2315
2324

65
102
170
200

102
100
81
78

Figure 10-28. Carbonate ion content of high-latitude surface water

as a function of atmospheric C02 partial pressure.
assumed to be at chemical equilibrium with the air.

The water is

140 ....- - - - - - -....- -.....- -.....- -....- -....

~

T= 2.5°C
S= 34.95%0

120'"-Eo 100

[AL K] = 2290 ,ueq/kg:l

II rt')

o

80

U

60
300

500

700

P~~: (I0-6atm )
559

900

sediment.
Complete neutralization will require many "cycles"(hence many times 500 years).
It is for this reason that we said
early in the chapter that the time scale for neutralization of thefossil fuel CO 2 by calci te in the sediment will take several
thousand years.
This concept can be stated in another way which may make iteasier to grasp.
It is likely that man will consume over the next
century or two on the order of 15 x 10 16 moles of fossil fuel carbon.
Thus at a calcite dissolution rate of 5 x 10 13 moles/year(one tenth the current C02 production rate) it would take 3000
years to neutralize this CO 2 .
There is one big IF in all this. That IF has to do with thedependence of solution rate on carbonate ion concentration.
We
have assumed this dependence to be first order (i.e., the percentage increase in solution rate is equal to the percentage reduction
in carbonate ion content of deep water).
There is, however, evidence that this might not be the case.
Experiments conducted in
the laboratory yield a reaction rate which depends on the fourth
power of ~C03= (see figure 10-29).
If so then the solution rate
for water with a ~C03= of 20 lJm/kg would be 16 times (rather
than 2 times) that for water with a bC03= of 10 lJm/kg!
As theexperiments which yielded this dependence were carried out on
calcite particles suspended in a very rapidly stirred solution
they are not necessarily a valid analogue to what happens on thesea floor.
In the rapidly stirred laboratory flask the "stagnant"boundary film surrounding each g~ain is likely to be so thin thatit does not limit the rate of solution. Rather, the rate at whichions escape from the crystal surfaces controls the rate of solution under these condi tions .
The ques tion then is whether therate at which ions escape from the crystal surfaces also controls
the resaturation time for the waters in the pores of marine sediments.
If it does, then the solution rate will increase with increasing ~C03= much faster than we have assumed and the neutralization time will correspondingly diminish (as long as the resistance posed by the benthic boundary layer is not the limiting
resistance).
Two other barriers must be considered.
One is diffusion fromthe crystal face to the interior of the adjacent pore and theother is coatings of organic material over the calcite.
The importance of the former can be assessed by comparing the time required for an ion to diffuse the "width" of a pore with the time
for resaturation.
As long as the transit is far less than the resaturation time diffusion from crystal face to pore interior is
not the rate limiting step.
The time, T, for an ion to traverse apore by diffusion can be approximated as follows:
~2

10-37

T

D

where ~ is the pore width and D is the ion diffusi vi ty.
In calcite rich deep sea sediments the average pore width is on the order of 10 microns (1 x 10- 3 cm) and the diffusivity is on theorder of 4 x 10- 6 cm 2/sec.
Hence the time for an ion to traverse

560

100--.-........~....

-

>c
~
~
0
.......,.

w

10

I<l
0:

Z
0
I-

-

:::>

...J

0

en

1.0

(/)

-

0

0.1 ~--a.-A __~"""
0.1

1.0

[CO; ]SAT-[CO;][CO =]
3 SAT
Figure 10-29. Laboratory measurements of the rate of solution of
calci te grains suspended in rapidly stirred solu tions maintained
at various degrees of undersaturation.
As can be seen, a IO-foldchange in the degree of undersaturation leads to roughly a 10,000fold change in solution rate. These results were obtained by Keir
(479) while working in the laboratory of Berner at Yale University.

561

a pore should be on the order of one second.
As the pore water
resaturation time is on the order of 1 minute, this simple calculation suggests that the resaturation process is not limited by
the rate of diffusion of ions from crystal face to pore interior.
The extent to which the calcite in marine sediments is coated with an organic film is not known.
Howeve r, as the organic
content of these sediments is very low, we might be safe in assuming that bacteria have removed the organic sheathing which surrounds the platelets of calcite in living organisms.
In chapter 2 we mentioned the possibili ty that the rate of
calcite solution on the sea floor might be limited by the presenceof a "stagnant film" just above the sediment water interface.
Ifsuch a film constituted the major resistance to calcite solution,then first-order kinetics would be entirely appropriate.
Asstated in chapter 2, we do not yet have the information needed toevaluate the role of this resistance on the sea floor.
At thetime this book was being edited, the first experiment designed bySantschi and his colleagues at the Lamont-Doherty Geological Observatory to measure this resistance was completed.
Flat platesof alabaster (CaS04- 2H 2 0) were brought to the sea floor by thesubmersible ALVIN and allowed to dissolve under sea floor conditions for four days before recovery_ A stagnant film thickness of
110 microns was obtained. Further studies of this type will allow
the importance of this film in limiting the rate of fossil fuelinduced calcite solution to be defined.
One other process which must be understood if we are ever tomake quantitative predictions of the rate of excess calcite solution induced by fossil fuel C:O 2 is the rate at which organisms
churn the upper several centimeters of sediment.
Is this process
sufficiently rapid with respect to the rate of solution that thecalci te content of the bioturbated layer assumed to remain uniform, or will an insoluble residue build up at the sediment waterinterface~
Such a residue wquld impede the rate of solution.
The
distribution of the isotope 210 Pb has something important to ,tell
us in this regard.
As we learned in chapter 4 much of the 21 °Pb
produced by the decay of the 226Ra within the sea and added to theocean from the atmosphere reaches the sea floor before undergoingradiodecay.
These excess 210 pb atoms trace the bioturbation process.
Measurements of the concent ra tion of excess 2 1 °pb (over226 Ra ) as a function of depth in deep sea cores show that theseatoms are able to penetrate to a depth averaging about 2 cm intheir mean lifetime of 32 years (see figure 10-30).
If the bioturbation 'process is assumed to follow the rules of diffusion (alikely but unproven assumption) then the effective stirring depth
will vary as the square root of time.
Thus mixing to an average
depth of 4 cm would take 128 years and to a depth of 8 cm, 512
years.
As the apparent radiocarbon ages change by only a few hundred years from the top to the base of the bioturbated zone (figure 5-13) they are consistent with this 210Pb-based mixing scenario.
When compared to the projected rates of fossil fuel C:02 induced calcite solution these stirring rates prove rapid enough tomaintain a nearly homogeneous calcite content in the bioturbatedlayer.
Hence the worms living on the sea floor will do their partin promoting the transfer of fossil fuel CO 2 from air to sea!

562

Figure 10-30. Plots of excess 210 Pb activity versus depth in two
deep sea cores.
The core on the left was taken by Be rge r at
Scripps, and the 210 Pb analyses were made at Lamont-Doherty Geological Observatory (245). The core on the right was collected bydivers in t0e submersible ALVIN during a study of the Mid-Atlantic
Ridge and the 210 Pb analy ses were made by Nozaki and Cochran at
Yale (243). The solid curves are those expected if the 210 Pb werebeing moved downward by a mechanical stirring process which follows the mathematics of diffusion.
The required rate is about
one-tenth that for molecular diffusion of dissolved ions in sea
water.

563

SUMMARY

Over the next 200 years man will, through his energy and land
management practices, release an amount of CO 2 equal to somewhere
between 2 and 5 times that present in the atmosphere in the early
1800's.
Because the buildup of C02 in the atmosphere will modify
our planet's climate there is considerable interest in the rate at
which the sea will take up this man-made CO 2 . The transfer of C02
from the atmosphere to the sea is made possible by the C03= ion
dissolved in the sea and by the CaC0 3 in the sediments lining thesea floor.
Our knowledge of the rate at which this uptake process will
occur comes almost entirely from our knowledge of the distributionof tracers in the atmosphere-ocean system.
These tracers tell us
the rate of invasion of COz into the sea, they tell us of the rateof stirring in the sea, they tell us of the rate of dissolution of
CaC03 on the sea floor and of the rate of s ti rring of the uppe r
sediment column.
This knowledge allows us to calibrate models of
the fossil fuel C02 uptake process.
Fortunately, enough information is left over after the model has been constructed to allow
crosschecks on the model's reliability to be made.
While the model by necessity is far more simplified than one would like, thefact that it can adequately reproduce a very diverse set of tracer
distributions leads us to believe that the answers it gives withregard to the uptake of man-made CO z by the sea are reasonably
good.
The resul ts show that the uptake by the various reservoir s
are on very different time scales.
The surface mixed layer equilibrates with the atmosphere in about one year.
The waters of themain ocean thermocline equilibrate with the atmosphere on the timescale of several tens of years (see superproblem #10). The watersof the deep sea equilibrate with the atmosphere on time scales ofmany hundreds of years.
The calci te in marine sediments will
equilibrate with the atmosphere on the time scale of several thousands of years.
Finally, the excess Ca dissol ved in the sea as
the result of the calcite-C.Oz reaction will be removed on the timescale of many tens of thousands of years (see superproblem #10).
Hence the effects of the C02 added to the atmosphere by man overthe next two centuries will alter the ocean's chemistry even on
geologic time scales!
PROBLEMS

10-1

From .the information given in this chapter estimate what theatmospheric f).14 C value would have been in 1973 were no radiocarbon to have been produced by nuclear testing.

10-2

If at some time in the next hundred years the C02 pressureof the atmosphere reaches twice its present value by roughly
how much will the pH of the surface ocean have changed?
Do
this calculation using simplifying assumptions similar tothe constant PCO z x [C03=J approximation.

10-3

Assume that the annual use of fossil fuels rises at the rateof 2%/year until the year 2020 and then remains constant at
the 2020 value until the year 2060.
Based on the informa-564

tion given in this chapter estimate the atmospheric C02 partial pressure in the year 2060.
Also estimate the Keelingfraction at that time.
10-4

Assume that a decision is made by the global leaders of 2050
not to permit the C02 content of the atmosphere to rise any
further.
At this time the C.02 partial pressure has alreadyreached 600 x 10- 6 atmospheres.
The total fossil fuel CD 2
production up to that time is 8.3 x 10 16 moles.
The meanage of the fossil fuel CD 2 molecules at that time is 40
years.
How much production would this permi t during thenext year if the CO 2 pressure were to be maintained at 600 x
10- 6 atmospheres?
How much would be permi tted in 2080 if
this lid on CO 2 pressure were continued?

SUPERPROBLEM #10
Dieter Netherland, a geochemist, becomes interested in thelong term effects of fossil fuel C02 on ocean-atmosphere chemistry.
He realizes that even after all the fossil fuels have been
burned and the CO 2 produced in this way has been mixed throughout
the ocean and come to equilibrium wi th the calci te in the sediments, the chemistry of the oceans and the atmosphere will nothave returned to their preindustrial composition. The C02 content
of the atmosphere and the Ca content and EC02 content of the ocean
will be higher.
Netherland devises a hypothetical ocean to guidehis thinking in this matter.
His model is shown in figure 10-31.
The model's lithosphere consists of a mixture of three solid
compounds, CaSiO 3, CaC 03 and SiO 2.
Calci te and opal (bu t not
CaSiO 3) are precipitated from the waters of his model ocean.
These phases along with detrital CaSi0 3 (as the mineral wollastonite) and detrital 8i02 (as the mineral quartz) make up the sediment on hi-s model's sea floor.
The accumulation of new sedimentis matched by subsidence of the floor of the ocean.
The sinking
sediment eventually becomes buried to a depth at which the temperature is high enough to react the CaCO 3 and SiO 2 to form CaSiO 3
and CO 2.
(Any leftover opal is converted to quartz).
The wollastonite and quartz make up the model's metamorphic crust.
The
C02 (released by the breakdown of calcite) and the H20 (releasedby the breakdown of opal) are dri ven to the surface and become
part of the ocean-atmosphere system.
The metamorphic rock ispushed around and up to the surface of the model's continent.
Here it undergoes weathering.
C02 from the atmosphere reacts withthe wollastoni te to produce Ca ++, HCO 3-, and H48i04 ions.
The
weathering reaction is as follows:

The ions so formed are carried by the model rivers to the model
ocean.
Here they are reprecipitated as calcite and opal completing the grand geochemical cycle.
Netherland envisions a rather simple set of rules controlling the flow of material through his model ocean prior to its industrial revolution.
The tectonic processes in the earth set the

565

.

.

-

.

........................................................................
.
.. .. . .. .. .. .. . . .. . .. . .. .. .. .. .. .. .. .. .. .. .. .. .. .. .. . .. . .. .. .. .. .. .. .. .. .. .. . .. . .. .. . .. .. .. .. .. .. . . . .. .. .. .. .. .. ".. .. .. .. .. . .. . .. . .. .. .. .. .. .. . . .. . .. .. .. . .. .. .. .. . .. .. .. . .

::::::::::::::::::::::::::::::::::::A'.. 'M'.. O·.. S·.. 'p'. ·H·. ·E·.. R'.. E'.'::::::::::::::::::::::::::::::::.. . .. .. .. .. .. . .. .. .. . .. .. .. .. . .. .. . .. .. .. . .. .. .. .. . .. . .. .. .. .. .. . 'T'..
. ..
.... .. ... .. .... .. .... .. ..... .... .. .... .. ... .. ... .. ... .. ... . .... .. .... .. .... .. ... .. .... . .... .. .... ..

.,..

.. .

:...: : :·Ca...'++..+ 'H'.C·0 3- :::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::::.... .
. . . .. . .. .. . . .. .. .
.. . . . ..... .
.. .. ...... .... .. . ..

.... .. ...... .... .... ..

..

.:.:.:.:
H4 5 ·I 0 4·····:.:-..
..
.. .... .. ... .. .... .. .... +.....
.. .. . ..... ..... .... .. .. .. .. .. .. .. .. .. .. .. .. .... . .. .. .. . ... ...
..
. ...........
.. . . .. . .. .. ...
..........
.. ... . . .. .

.. .. .. .. .. .... .... .. ...... .. .. .... .... ...... .... .. .. .. .. .. .. .. ...... ...... ... .. .. .. .... .. ...... .... .... .... .. .. .. ..
0•••••••

0.... ....

•••••••••••••••••..

0.:';.:.:':-:':':':':-:':':':':':.:.:-:':':':"..:

-

OCEAN

CaSi0 3 +Si02 CO2

LITHOSPHERE

Figure 10-31. Dieter Netherland's simplified earth cycling modelof the earth's geochemical cycles (see Superproblem #10).

566

rate at which the crust moves from sea floor to continent.
In sodoing it determines the rate at which C02 gas percolates out of
the lithosphere into the ocean-atmosphere system.
As this C02 canonly be precipitated onto the sea floor (as CaC03) as fast as
CaSi0 3 is weathered on the continents the weathering rate must at
steady state match the C02 production rate by metamorphism.
Netherland envisions that the weathering rate depends on the CO 2
partial pressure in the atmosphere - the higher the pressure thethe higher the weathering rate.
Hence the partial pressure of C02
in the atmosphere prior to industrialization had achieved thatlevel where the weathering rate matches the rate at which C02 is
produced by metamorphism.
Finally, he assumed that his model
ocean remains close to saturation with respect to both opal andcalcite. This fixes the chemistry of his model's ocean. The concentrations of its five ions (i.e., C03=' HC03-, C02, Ca++and H4Si04) are fixed by the following restrictions:1) charge neutrality
2[Ca++J + [HC0 3-J + 2[C03=J =°2) weathering rate balances metamorphism rate
[C02] = constant3) opal saturation[H 4Si0 4 J = constant4) calcite saturation[Ca++J [C03=J = constant5) equilibrium among the carbon species[HC0 3 -J 2

constant[ C03=J[ C0 2J
Netherland now assigns
state system:
Vocean1.0

Xsome numerical values to his

10 18 m3

4.0 x 10 14 m3
Vo/A o = 2.5 x 10 3m[CO 2J1 x 10- 2 moles/m 3[HC0 3 -J

200 x 10- 2 moles/m 3[C0 3 =]

30 x 10- 2 moles/m 3[Ca++J

130 x 10- 2 moles/m 3[H 4SiO 4J

60 x 10- 2 moles/m 3567

steady

CO 2

M

atm

2.3 x 10 16 moles= 2.3 x 10 13 mo1es/yr

These concentrations of course fix the constants for restrictions
2 through 5.
Now Netherland is ready to subject his model to an industrial revolution.
To do so he injects 2.3 x 10 17 moles of CD2 into
the ocean atmosphere system.
He allows this C02 to mix throughthe ocean and attack the CaC03 in the ocean's sediments until calcite saturation is reestablished.
At this point he computes thefollowing composition for its ocean (and its atmosphere).
[ C0 2]= 1.8 x 10- 2 moles/m 3[HC0 3-J

248 x 10- 2 moles/m 3[C03=J

26 x 10- 2 moles/m 3[Ca++J

150 x 10- 2 moles/m 3[H 4 Si0 4 ]60 x 10- 2 moles/m 3

Do these values satisfy the restrictions 1, 3, 4 and 5?
Now Netherland is ready to consider the long-term future ofhis model ocean.
Even after equilibration with the sea and its
sedimentary calcite, the excess 'industrial C02 has left Netherland's model ocean with an atmospheric C02 content about 1.8 times
higher than its preindustrial value.
This higher atmospheric C02
content would of course lead to a higher continental weathering
rate.
Netperland envisions two scenarios. First, he assumes that
continental weathering rate is proportional to the atmospheric CD2pressure.
If so, how long will it take for the atmospheric C02
content to drop halfway back to its preindustrial value?
Netherland is concerned wi th this resul t, however.
He isaware that the decay of terrestrial plant material leads to CO 2pressures in soils averaging about 10 times the steady state atmospheric C02 pressure.
He decides to repeat the calculation assuming that the soil CO 2 pressure controls the weathering rate andthat this pressure increases only the same increment as the atmospheric COo2 pressure.
To calculate the new soil CO 2 pressure heuses the following equation:
soil
P. I. A. ( a tm
pe02 = 10 Pc a 2 + PC a 2-

P . I • A e)
P. I • A•
PC a 2= 9 Pc a 2

P.I.A.
where pea 2
is the steady state preindus trial atmospheric CO 2pressure.
Under these circumstances how long would it take forthe atmospheric CO 2 content to drop halfway to its preindustrial
value?

568

f< elelJ-enGe~
INTRODUCTION
The references are arranged according to a subject outline.
In each sub-category the references are placed in order of publication.
For each reference a brief summary of its bearing on the
mission of this book is given.
The numbers (1 through 737) areused to designate specific references in the text, table andfigures.
SlBJECT OUTLINE

I.

PAPERS DEALING WITH RADIOTRACERS (#1 to #264)
A. PAPERS DEALING WITH RADIOTRACERS OF WATER MOTION (#1 to
#111)

1.

PAPERS DEALING WITH THE DISTRIBUTION OF NATURAL RADIOCARBON IN THE OCEAN-ATMOSPHERE SYSTEM (#1 to #26)

2.

PAPERS DEALING WITH MODELING OF THE DISTRIB UTION OF
RADIOCARBON IN THE OCEAN-ATMOSPHERE SYSTEM (#27 to #41)

3.

PAPERS DEALING WITH THE DISTRIBUTION OF TRITIUM IN THE
OCEAN (#42 to #56)

4. -- PAPERS DEALING WITH THE INTERPRETATION AND MODELING OF

THE ADDITION OF TRITIUM TO THE SEA AND ITS PENETRATION
INTO THE BODY OF THE SEA (#57 to #70)

B.

5.

PAPERS DEALING WITH THE JOINT INTERPRETATION OF THE
DISTRIBUTIONS OF BOMB RADIOCARBON AND TRITIUM IN THE
OCEAN (#71 to #75)

6"

PAPERS DEALING WITH THE DISTRIBUTION OF EXCESS 222 Rn
IN THE DEEP SEA (#76 to #86)

7•

PAPERS DEALING WITH THE DISTRIBUTION OF 228Ra IN THE
SEA (# 'd 7 to # 96 )'d.

PAPERS DEALING WITH THE DISTRIBUTIONS OF
85 Kr and 39 Ar IN THE SEA (#97 to #111)

90 Sr

,

137 Cs,

PAPERS DEALING WITH RADIOTRACERS OF PARTICLE MOVEMENT AND
THE SCAVENGING OF REACTIVE METALS (#112 to #170)
1.

PAPERS DEALING WITH THE DI;STRIB UTION OF
OCEAN (#112 to #121)

569

226

Ra IN THE

2.

PAPERS DEALING WITH THE DISTRIBUTION OF 234Th 228 Th
230 Th AND 231 Pa IN THE OCEAN (#122 to #138)',

3.

PAPERS DEALING WITH THE DISTRIB UTION OF 21 0po AND
IN THE SEA (#139 to #161)

4.

PAPERS DEALING WITH THE INPUT OF AND DISTRIBUTION OF
PLUTONIUM WITHIN THE SEA (#162 to #165)

5.

PAP ERS DEALING WITH THE DISTRIB UTION OF 14 C IN PARTIC ULATE DEBRIS AND ORGANISMS (#166 to #167)

6.

MODELS FOR PARTICULATE SCAVENGING IN THE OCEAN (#168 to#170)

2 1 0p b

C. PAPERS DEALING WITH RADIO TRACERS OF GAS EXCHANGE (#171 to
#194)

D.

II.

1.

PAPERS DEALING WITH THE RADON METHOD (#171 to #177)

2.

PAPERS DEALING WITH THE 3H - 3He METHOD (#178 to #179)

3•

PAPERS DEALING WITH THE RADIOCARBON METHOD (#180 to
#184)

4.

PAPERS DEALING WITH LABORATORY MEASUREMENTS OF GAS EXCHANGE RATES (#185 to #188)

5.

THEORETICAL TREATMENTS OF VARIOUS ASPECTS OF THE GASEXCHANGE PROCESS (#189 to #194)

PAPERS DEALING WITH RADIOISOTOPES IN OCEAN SEDIMENTS (#195
#264)

to-

1.

PAPERS DEALING WITH RADIOCARBON MEASUREMENTS IN DEEPSEA SEDIMENTS (#195 to #199).

2.

PAPERS DEALING WITH 23 QTh AND 231 pa DATA IN DEEP SEA
SEDIMENTS AND MANGANESE NODULES (#200 to #239)

3.

PAPERS DEALING WITH 210 Pb AND OTHER BIOTURBATIONTRACERS IN DEEP SEA SEDIMENTS (#240 to #247)

4.

PAPERS DEALING WITH THE DISTRIB UTIONS OF 26 A1 , 32Si AND
lOBe IN SEDIMENTS AND MANGANESE NODULES (#248 to #264)

PAPERS DEALING WITH STABLE ISOTOPE TRACERS (#265 to #345)
A.

PAPERS DEALING WITH STABLE ISOTOPE TRACERS OF WATER MOTION(#265 to #2?34)
1.

PAPERS DEALING WITH THE DISTRIBUTION OF THE
RATIO IN THE SEA (#265 to #281)

570

3He /4 He

2.

B.

C.

PAPERS DEALING WITH STABLE ISOTOPE TRACERS OF BIOLOGICALPROCESSES IN THE SEA (#285 to #317)
1..

PAPERS DEALING WITH THE 13C/12C RATIO IN DISSOLVED
OCEANIC CARBON AND IN MARINE ORGANISMS (#285 to #305)

2.

PAPERS DEALING WITH THE 180/160 RATIOS IN ATMOSPHERIC
02 AND IN THE 02 DISSOLVED IN THE SEA (#306 to #317)

STABLE ISOTOPE TRACERS IN SEDIMENTS AND MANGANESE NODULES(#318 to #345)
1.
2.

III.

PAP ERS DEALING WITH THE DISTRIB UTIONS OF D/H AND OF
180/160 RATIOS IN THE OCEAN (#282 to #284)

PAPERS DEALING WITH THE 180/160 RATIOS IN CaC03 FROMDEEP SEA SEDIMENTS (#318 to #326)
PAPERS DEALING WITH THE 13 C/ 12 C RECORD FOUND IN
FORAMINIFERA FROM DEEP SEA SEDIMENTS (#327 to #331)

3.

PAPERS DEALING WITH 143Nd /144 Nd RATIOS IN SEA WATER,
MARINE SEDIMENTS AND MANGANESE NODULES (#332 to #340)

4.

PAPERS DEALING WITH LEAD ISOTOPE RATIOS IN DEEP SEA
SEDIMENTS AND MANGANESE NODULES (#341 to #345)

PAPERS DEALING WITH CHEMICAL TRACERS (#346 to #456)
A.

PAPERS DEALING WITH CHEMICAL TRACERS OF BIOLOGICAL CYCLES(#346 to #433)
1. -- PAP ERS DEALING WITH THE DISTRIB UTIONS OF NO 3, P04 AND
DISSOLVED OXYGEN IN THE SEA (#346 to #364)
2.

PAPERS DEALING WITH THE DISTRIB UTIONS OF H4S10 4, Ge AN D
Ba IN THE SEA (#365 to #378)

3.

PAPERS DEALING WITH THE DISTRIBUTIONS OF rC02, C03=
AND ALKALINITY IN THE SEA (#379 to #409)

4. ',PAPERS DEALING WITH THE DISTRIBUTION OF NITROGEN'SPECIES (OTHER THAN NITRATE IN THE SEA (#410 to #421)

5.

B.

PAPERS BASED ON THE GEOSECS DATA SET DEALING WITH THE
RELATIONSHIP BETWEEN POTENTIAL TEMPERATURE, SALINITY
AND THE DISTRIB UTIONS OF SEVERAL OF THE NUTRIENTCONSTITUENTS (#422 to #433)

CHEMICAL TRACERS OF PARTICLE-WATER INTERACTIONS OCCURRING
IN THE SEA (#434 to #456)
1.

PAPERS DEALING WITH THE DISTRIBUTION OF LEAD IN THE SEA(#434 to #436)

571

IV.

2.

PAPERS DEALING WITH THE DISTRIBUTION OF CADMIUM IN THE
SEA (#437 to #438)

3•

PAPERS DEALING WITH THE DISTRIBUTION OF COPPER IN THE
SEA (#439 to #441)

4.

PAPERS DEALING WITH THE DISTRIBUTION OF ZINC AND NICKELIN THE SEA #442 to #443)

5·

PAPERS DEALING WITH THE DISTRIB UTION OF MANGANESE IN
SEA WATER (#444 to #449)

6.

PAPERS DEALING WITH THE DISTRIBUTION OF SELENIUM IN
SEA WATER (#450 to #451)

7•

PAPERS DEALING WITH SEVERAL METALS (#452 to #456)

PAPERS DEALING WITH BIOGENIC MATERIAL (#457 to #543)
A. PAPERS DEALING WITH BIOGENIC MATERIAL IN THE SEDIMENTS(#457 to #531)
1.

PAPERS DEALING WITH THE SOLUBILITY OF CALCITE AND
ARAGONITE (#457 to #468)

2.

PAPERS DEALING WITH CaC0 3 SATUROMETRY (#469 to #472)

3.

PAPERS DEALING WITH THE RESULTS OF CALCITE AND
ARAGONITE IN SITU DISSOLUTION EXPOSURES (#473 to #476)

4.

LABORATORY STUDIES OF THE DEPENDENCE OF THE RATES OF- CALCITE AND ARAGONITE DISSOLUTION RATE ON THE CARBONATE- ION CONTENT (HENCE pH) OF THE WATER (#477 to #479)

5.

PAPERS DEALING WITH THE DISTRIBUTIONS OF CALCITE AND
ARAGONITE IN RECENT SEDIMENTS (#480 to #495)

6.

PAPERS DEALING WITH GLACIAL TO INTERGLACIAL CHANGES AND
ARAGONITE DISSOLUTION (#496 to #512)

7 •. MODELS OF THE DISSOLUTION OF CALCITE ON THE SEA FLOOR(#513 to #519)

B.

8.

PAPERS DEALING WITH THE DISTRIBUTION OF OPAL IN MARINESEDIMENTS (#520 to #523)

9.

PAPERS CONTAINING PORE WATER PROFILES OF DISSOLVEDSILICATE (#524 to #529)

10.

PAPERS CONTAINING DATA REGARDING THE DISTRIBUTION OF
ORGANIC MATERIAL IN SEDIMENTS (#530 to #531)

PAPERS DEALING WITH BIOGENIC MATERIAL IN THE WATER COLUMN(#532 to #543)

572

V.

1.

PAPERS DEALING WITH SUSPENDED MATTER IN THE WATERCOLUMN (#532 to #536)

2.

PAPERS DEALING WITH MATERIAL CAUGHT IN SEDIMENTTRAPS (#537 to #541)

3.

PAPERS DEALING WITH MODES OF TRANSPORT OF PARTICULATE
MATTER THROUGH THE WATER CDLUMN (#542 to #543)

PAPERS DEALING WITH CONTROLS ON OCEAN CHEMISTRY (#544 to #547)

VI.

A.

PAPERS CONCERNED WITH THERMODYNAMIC CONTROLS ON OCEANCHEMISTRY (#544 to #546)

B.

PAPERS CONCERNED WITH KINETIC CONTROLS ON OCEAN CHEMISTRY(#547)

PAPERS CONCERNED WITH MAN-MADE PERTURBATIONS IN THE EARTH'S
CARBON CYCLE (#548 to #599)
A.

PAPERS DEALING WITH THE PRODUCTION OF FOSSIL FUEL C02, ITS
BUILD-UP IN THE ATMOSPHERE, AND ITS DISTRIBUTION WITHINTHE ATMOSPHERE (#548 to #562) .

B.

PAPERS DEALING WITH CHANGES IN THE LIVING BIOMASS AND SOIL
CARBON RESERVOIRS (#563 to #574)

C.

PAPERS DEALING WITH THE UPTAKE OF ANTHROPOGENIC C02 BY THE
SEA (#575 to #597)

D.

PAPERS DEALING WITH THE REAC~ION 'BETWEEN ANTHROPOGENIC C02
CaC0 3 IN MARINE SEDIMENTS (#598 to #599).

A~D

VII.

PAPERS DEALING WITH THE GLACIAL TO INTERGLACIAL CHANGE IN
VARIOUS OCEAN PROPERTIES (#600 to #624)
A.

PAPERS DEALING WITH THE RECONSTRUCTION OF SEA SURFACETEMPERATURE (#600 to #610)

B.

PAPERS DEALING WITH THE RECONSTRUC~ION OF ICE VOLUME AND
HENCE OCEAN SALINITY (#611 to #617)

C.

PAPERS DEALING WITH CHEMICAL CHANGES IN THE OCEAN AND
ATMOSPHERE (#618 to #624)

VIII.

PAPERS DEALING WITH RATES OF INPUT, RATES OF LOSS, ANDRESIDENCE TIMES OF VARIOUS CDNSTITUENTS IN THE SEA (#625 to
#670)
A.

PAPERS DEALING WITH RIVER INPUTS TO THE SEA (#625 to
#633)

B.

PAPERS DEALING WITH THE REMOVAL RATES OF CDNSTITUENTS TO
DEEP SEA SEDIMENTS (#634 to #635)

573

C.

PAPERS DEALING WITH INTERACTIONS BETWEEN SEA WATER AND
RIDGE CREST BASALTS (#636 to #643)

D.

PAPERS DEALING WITH ATMOSPHERIC AEROSOLS AND DUST (#644 to
#652)

E.

PAPERS DEALING WITH GAS SOLUBILITY IN SEA WATER (#653 to
#660)

F.

PAPERS DEALING WITH MOLECULAR DIFFUSIVITY OF GASES IN SEAWATER (#661 to #666)

G.

PAPERS DEALING WITH THE EVOLUTION OF OCEAN CHEMISTRY(#667 to #670)

IX.

BOOKS ON RELATED SUBJECTS (#671 to #684)

X.

ATLASES CONTAINING CHEMICAL DATA (#684a to #685g)

XI.

MISCELLANEOUS PAPERS AND LATE ENTRIES (#686 to #737)
ANNOTATED REFERENCE LISTI.

PAPERS DEALING WITH RADIOTRACERS (II to 1264)
A.

PAPERS DEALING WITH RADIOTRACERS OF WATER MOTION (#1 toIll)
1.

PAPERS DEALING WITH THE DISTRIBUTION OF NATURAL RADIOCARBON IN THE OCEAN-ATMOSPHERE SYSTEM (11 to #26)

1.

Suess,~H.E.
1955. Radiocarbon concentration in modern wood.
Science, V122, p. 415-417.
The first measurements showing the decline in the atmospheric 14C/ C ratio due to the build up of fossil fuel C02
in the atmosphere and the first demonstration that the magnitude of this decline could be used as a measure of theextent of mixing between atmosphere, biosphere and ocean
carbon on the time scale of several decades.

2.

Vries,·.H. de. 1958. Variations in concentration of radiocarbon with time and location on the earth: Koninkl Nederlandse
Akad Wetensch. Proc. sec. B. V61, p. 1-9.
The first demonstration that the atmosphere has undergone
short-term variations in 14C/C ratio due to natural causes.

3.

Fergusson, G.J. 1958. Reduction of atmospheric radiocarbon
concentration by fossil fuel carbon dioxide and the mean lifeof carbon dioxide content in the atmosphere. Proc. Roy. Soc.
London A, V243, p. 561-574.
A follow up to the work of Suess involVing accurate 14 C/ C
ratio measurements on samples from southern hemispheretrees covering the period 1890 to 1955.

574

4.

Broecker, W.S., R. Gerard, M. Ewing, and B.C.. Heezen. 1960.
Natural Radiocarbon in the Atlantic Ocean. Jour. of Geophys.
Res. V65, p. 2903-2931.
Radiocarbon measurements on samples from throughout the
Atlantic Ocean collected before significant amounts ofbomb testing 14 C had reached the oceans.

5.

Bien, G.S., N.W. Rakestraw and H.E. Suess. 1960. Radiocarbon
concentration in Pacific Ocean water, Tellus V12, p. 436-443.
Radiocarbon measurements on samples from the Pacific Ocean
collected before significant amounts of bomb testing 14 C
had reached the ocean.

6.

Burling, R.W., and D.M. Garner. 1960. A section of 14C activities of sea water between 90 S and 66°s in the southwest
Pacific Ocean, New Zealand J. Geol. Research V2, p. 2903.
Radiocarbon measurements on samples from the Pacific Ocean
collected before significant amounts of bomb testing 14C
had reached the ocean.

7.

Nydal, R. 1963. Increase in radiocarbon from the most recent
series of thermonuclear tests. Nature, V6, p.212-214.
Measurements on atmospheric C02 showing the increase in the
14C/C ratio resulting from nuclear testing.

e.

Broecker, W.S. 1963.
14C/12 C ratios in surface ocean water.
Nat. Acad. Sci., Nat. Res. Counc. Publ. 1075, p. 138-149.
A summary of radiocarbon measurements made on surface ocean
water and shell samples collected prior to the onset of thecontamination of the ocean with nuclear radiocarbon.

9.

Munnicn, K.O., and W. Roether. 1967. Transfer of bomb 14C andtritium--from the atmosphere to the ocean: Internal mixing of
the ocean on the basis of tritium and 14C profiles, in: Symposium on Radioactive Dating and Methods of Low Level Counting.
International Atomic Energy Agency UN DOC SM-87/22, p. 93-103.
Measurements of bomb radiocarbon in atmospheric C02 overthe period 1959 to 1966 and in the North Atlantic Ocean
samples collected in 1966.

10. Nydal, R. 1968. Further investigation of the transfer of radiocarbon tn nature, Jour. Geophys. Res. V73, p. 3617-3635.
This paper describes the transfer of bomb-produced radiocarbon from stratosphere to troposphere and from theatmosphere to the mixed layer of the ocean.
11. Rafter, T.A. 1968. Carbon-14 measurements in the South Pacificand Antarctic Oceans. New Zealand Jour. Sci., VII, p. 551-589.
Measurements of the bomb 14 C in surface waters in the South
Pacific over the time interval 1954 to 1968.
12. Broecker, W.S., and R. Gerard. 1969. Natural radiocarbon in
the Mediterranean Sea. Limnology and Oceanography, V14, p.
883-888.

575

The distribution of radiocarbon in the Mediterranean Sea
prior to nuclear testing. Estimates of the ventilation
times of the deep waters in the eastern and western basinsare given.
13. Linick, T.W. 1978. La Jolla Measurements of Radiocarbon in the
Oceans. Radio carbon, V20, p. 333-359.
Radiocarbon measurements on samples from the Pacific andIndian Oceans collected after the onset of the build-up ofbomb radiocarbon in the oceans.
14. Nozaki, Y., D.M. Rye, K.K. Turekian and R.E. Dodge. 1978. A
200 year record of carbon-13 and carbon-14 variations in a
Bermuda coral. Geophys. Research Letters, V5, p. 825-828.
The first paper on the radiocarbon record in corals. For
reasons which are not clear this record is not entirely
representative of the surrounding ocean.
15. Druffel, E.M. and T.W. Linick. 1978. Radiocarbon in annual
coral rings of Florida. Geo~hYs. Res. Letters, V5, p. 913-916.
A reconstruction of the 4C/ C ratio in surface ocean waterin the Florida Straits from 1800 to 1974 based on measurements in growth ring dated corals.
15a.Gillespie, R. and H.A. Polach. 1979. The suitability ofmarine shells for radiocarbon dating of Australian prehistory.
Radiocarbon Dating, edited by R. Berger and H. Suess, University of California Press, Los Angeles, p. 404-421.
14C/C ratios on shells collected before 1950 along the east
coast of Australia yielding estimates of the prenuclear
14C/C ratios in surface water carbon from the western
Pacific. Also given are the results on shells collected in
19r3 showing the bomb 14 C build-up in this area.
16. Druffel, E.M. 1980. Radiocarbon in annual coral rings of
Belize and Florida. Radiocarbon, V22, p. 363-371.
Reconstruction of the 14C/C ratio in surface ocean water ofthe eastern Caribbean Sea from 1890 to 1977 based on
measurements in growth ring dated corals.
17. Druffel, E.M. 1980. Radiocarbon in annual coral rings of the
Pacific and Atlantic Oceans. Ph.D. thesis, Chemistry Dept.,
University of California, San Diego.
Radiocarbon measurements of corals from the Florida
Straits, from the Gulf of Honduras and from the Galapagos
Islands in the eastern equatorial Pacific.

Ie. Stuiver, M.

1980. 14C distribution in the Atlantic Ocean.
Jour. Geophys. Res. V85, p. 2711-2718.
Description of the distribution of bomb-produced
radiocarbon in the thermocline of the Atlantic Ocean.

19. Stuiver, M. and P.D. Quay. 1980. Changes in atmospheric 14C
attributed to a variable sun. Science, V207, p. 11-19.

576

Ultra-high accuracy measurements on tree ring samples of
the 14C/ C record for the period 1000 to 1900 AD. The
authors show that the variations are correlated with thesun spot, auroral and magnetic index records. Changes inthe sun's magnetic field have modulated the flux of c~~micrays to the earth's atmosphere and hence the rate ofCproduction.

20. Roether, W., K.O. Munnich, B. Ribbot, and J.L. Sarmiento.
1980. A transatlantic 14C-section near 40 o N. "Meteor"
Forsh.-Ergelenisse Reihe A, V22, p. 57-69.
Radiocarbon measurements on a track from Lisbon to Nova
Scotia run by the West German research vessel "Meteor" in
1971. These measurements nicely complement those of the
GEOSECS survey in the Atlantic Ocean.
21. Stuiver, M. and H.G. Ostlund. 1980. GEOSECS Atlantic Radiocarbon. Radiocarbon, V22, p. 1-24.
Radiocarbon measurements on samples collected as part of
the GEOSECS survey of the Atlantic Ocean during the period
July 1972 to April 1973.

22. Ostlund, H.G. and M. Stuivero
bon. V22, p. 25-53.

1980.

GEOSECS Pacific Radiocar-

Radiocarbon measurements on samples collected as part of
the GEOSECS survey of the Pacific Ocean during the period
September 1973 to June 1974.

1980. GEOSECS Indian Ocean Radiocarbon and Tritium Measurements (Miami). Tritium Laboratory Data Report #9, Rosenstiel School of Marine andAtmospheric Science, University of Miami.
Report containing the radiocarbon data obtained on samples
collected during the period Jan. 1978 to Feb. 1978 in thewestern Indian Ocean as part of the GEOSECS program.

23. Ostlund, H.G., R. Oleson, and R. Brescher.

1980. Transfer of
4C_ to the ocean surface" Radiocarbon, V22, p. 626-635.
Measurements of the bomb Ilt C_ content of atmospheric CO 2 andsurface ocean water over the period 1966 to 1978 from awide variety of localities in both hemispheres.

24. Nydal\ R., K. Lovseth and F.R. Skogseth.
bomb

Munnich~ and W. Weiss.
1980. The effect ofanthropogenic CO 2 and 1 C sources in the distribution of life in
the atmosphere. Radiocarbon, V22, p. 379-391.
lite measurements on weekly samples of atmospheric C02 andhydrocarbon to detect fossil fuel C02 dilution and nuclear
power plant life addition.

25. Levin, I., K.O.

1981. Atmospheric lite changes resulting from fossil fuel C02 release and cosmic ray flux variability. Earth and Planet. Sci. Lett., V53, p. 349-362.
The definitive set of measurements documenting the magnitude of the Suess effect and the shape of this decline.
Also an attempt is made to separate the contributions offossil fuel e02 and cosmic ray variations.

26. Stuiver, M. and P.D. Quay.

577

2.

27.

PAPERS DEALING WITH MODELING OF THE DISTRffiUTION OF
RADIOCARBON IN THE CX:EAN-ATMOSPHERE SYSTEM (#27 to 141)

C~raig,H. 1957 • The natural distribution of radiocarbon andthe exchange time of carbon dioxide between the atmosphere andsea. Tellus, V9, p. 1-17.
One of the Tellus triad of papers discussing theimplications of the distribution of 14C~ between theatmosphere and oceans. This paper lays out the principles
of box modeling in a very elegant way and thus provides
the foundation for the wide use of this approach to ocean
modeling.

2tl. Suess, H.E. and R. Revelle. 1957. Carbon dioxide exchange
between the atmosphere and ocean and the question of anincrease of atmospheric C02 during the past decades. Tellus,
V9, p. 18-27.
The second of the Tellus triad of pa~ers discussing theimplications of the distribution of 4C between the atmosphere and oceans. Using a one-box ocean model theseauthors made the first attempt to estimate the rate of uptake of fossil fuel C02 by the oceans.
29. Arnold, J.R. and E.C. Anderson. 1957. The distribution ofradiocarbon in nature. Tellus, V9, p. 28-32.
The third of the Tellus triad of papers discussing theimplications of the distribution of 14 C between theatmosphere and oceans.
30. Bolin, B. and E. Eriksson. 1959. Changes in the carbon
dioxide content of the atmosphere and sea due to fossil fuelcombu~tlon.
Rossby Memorial Volume, Rockefeller Inst. Press,
New York, p. 130-142.
The first application of the two-box ocean model to theuptake of fossil fuel C02 by the ocean.
31. Munk, W.H. 1966. Abyssal recipes. Deep-Sea Res. V13, p.
707-730.
The first attempt to calibrate a one-dimensional advection-diffusion model using radiocarbon data.
32. Keeling, C.D. and B. Bolin. 1967. The simultaneous use ofchemical tracers in oceanic studies 1. Tellus, V19, p. 5665Bl. Keeling, C.D. and B. Bolin. 1968. The simultaneous useof chemical tracers in oceanic studies 2. Tellus, V20,
p.17-54.
A very elaborate multibox ocean model designed to include
both the transport of ca~bon with the wate~ and withparticulate matter falling from surface water to the interior of the sea.
33. Craig, H. 1969. Abyssal Carbon and Radiocarbon in the
Pacific. Jour. of Geophys. Res. V74, p. 5491-5506.
A more rigorous attempt to calibrate a one-dimensional578

advection diffusion model using radiocarbon and carbon
data. Done with the characteristic Craig style, thispaper paved the way for further application of this typeof model by marine geochemists.
34. Broecker, W.S. and Y.-H. Li. 1970. Interchange of water
between the major oceans. Jour. of Geophys. Res. V75, p.
3545-3552.
A three-box model of the world ocean designed to evaluate the rate of exchange of water between the three majorwater types in the ocean: warm surface water, North
Atlantic deep water and Pacific and Indian deep water.
35. Stuiver, M. 1976. The 14C distribution in west Atlantic
abyssal waters. Earth and Planet. Sci. Lett. V32, p. 322-330.
A discussion of the distribution of radiocarbon with latitude in the deep waters of the western Atlantic based on
the GEOSECS data set.
36. Broecker, W.S., T.-H. Peng and M. Stuiver. 1978. An estimateof the upwelling rate in the equatorial Atlantic based on the
distribution of bomb radiocarbon. Jour. of Geophys. Res. V83,
p.6179-6186.
The deficiency in the water column inventory of bomb-produced radiocarbon in the equatorial ocean is explained by
the flushing action of upwelled water. In the paper ameans is presented to determine the flux of upwelledwater based on magnitude of this inventory anomaly.
37. Broecker, W.S. 1979. A revised estimate for the radiocarbon
age of North Atlantic deep water. Jour. of Geophys. Res. V84,
p. 3218-3226.
Much of the variation in the radiocarbon concentration in
deep water carbon from the Atlantic is the result of mixing of 14C-rich deep waters originating in the-northern
Atlantic with 14C_poor deep waters originating in thesouthern Atlantic. In this paper an attempt is made toseparate the contributions of miXing and radiodecay
within the Atlantic and to use the latter to estimate therate of ventilation of the deep Atlantic.
Bjorkstrom, A. 1979. A model of C02 interaction between
atmosphere, ocean, land biota. In: Global carbon cycle, ed. by
B. Bolin, Wiley & Sons, New York, p. 403-457.
A five-box ocean made up of the warm and cold surface
ocean, the upper and lower main thermocline and the deepsea is used to describe the transport of carbon in theworld ocean.
39. Broecker W.S. and T.-H. Pengo 1980. Seasonal variability
in the 1~C/12C ratio for surface ocean water. Geophys. Res.
Lett. V7, p. 1020-1022.
In the temperate zones of the oceans the surface mixed
layer thickens greatly during winter months. This leads579

to a seasonality in the bomb
surface water.

14 C

content of the carbon in

40. Quay, P.D. and M. Stulver. 1980. Vertical advection-diffusion
rates in the oceanic thermocline determined from 14C distribubutions. Radiocarbon, V22, p. 607-625.
Use of the one-dimensional advection-diffusion equation toestimate the apparent rates of vertical advection anddiffusion at individual GEOSECS stations in the Atlantic
Ocean.
41. Rooth, C.G.H., R.A. Fine and H.G. Ostlund. 1980. Abyssal mixing and circulation in the north Pacific from radiocarbon data.
Technical Report, Univ. of Miami, TR 80-2.

3. PAPERS DEALING WITH THE DISTRmUTION OF TRITIUM IN THE
OCEAN (142 to 156)
42. Buttlar, H.V. and W.F. Libby. 1955. Natural distribution ofcosmic ray produced tritium. Jour. of Inorg. Chern., VI, p.
75-95.
The discovery of natural tritium.
43. Giletti, B.J., F. Bazan, and J.L. Kulp. 1958. The geochemistry of tritium, EOS Trans. AGU, V39, p. 807-818.
Measurements of the tritium content of ocean water samples
collected prior to the onset of nuclear contamination.
Because of the very low tritium content of these samples
and the rather large blank correction applied, the validityof these measurements remains an open question.
44. Houtermans, J. 1~65. Tritium in surface water of the North
Pacific-Ocean. 6th Carbon-14 and Tritium Dating Conference,
CONF 650652, p. 565-572.
The distribution of bomb-produced tritium in the Pacific
Ocean.
45. Dockins, K.O., A.E. Bainbridge, J.C. Houtermans and H.E. Suess.
1967. Tritium in the mixed layer of the North Pacific Ocean.
In: Radiocative Dating and Methods of Low-Level Counting. Internatipnal Atomic Energy Agency, Vienna, p. 129-142.
Tritium measurements on surface water samples collected at
12 stations in the north and south Pacific over the period
1959 to 1966. Also given are measurements on surface water
samples collected along north-south traverses in the
Pacific Ocean during the same period.
46. Munnich, K.O. and W. Roether. 1967. Transfer of bomb- 14 C andtritium from the atmosphere to the oceans: Internal mixing of
the ocean on the basis of tritium and 14 C profiles. In: Radioactive Dating and Low-Level Counting. International Atomic
Energy Agency, Vienna, p. 93-104.
North-south traverses in the North Atlantic of the tritiumcontent of surface water taken during the period 1965-1966.

580

Also given are vertical profiles of tritium at stations
along these traverses.
47. Ostlund, H.G., M.O. Hinkle and C. Rooth. 1969. Tritium in theequatorial Atlantic Current System. J. Geophys. Res. V74,
p. 4535-4543.
Measurement of the vertical distribution of tritium atseveral equatorial Atlantic stations occupied in 1964.
48. Roether, W. 1974. The tritium and 14C profiles at the GEOSECSI (1969) and GOGO I (1971) North Pacific station. Earth &Planet. Sci. Lett. V23, p. 108-115.
Tritium and 14 C measurements at 28.5°N, 121.6°w in 1969
and 1971.
49. Ostlund, H.G., R. Brescher and W.H. Peterson. 1974. Ocean
tritium profiles 1965-1972. Tritium Laboratory Data Report#3. Hosenstiel School of Marine and Atmospheric Sciences,
University of Miami.
Tritium measurements from the north and equatorial
Atlantic on samples collected during the period 1965 to1972.
50. Michel, R.L. and H.E. Suess. 1975. Bomb tritium in the
Pacific Ocean. Jour. of Geophysical Res. vao, p. 4139-4152.
Extensive measurements of the distribution of tritium in
the Pacific Ocean over the period 1959 to 1973.
51. Ostlund, H.G., H.G. Dorsey and R. Brescher. 1976. GEOSECS
Atlantic Radio carbon and Tritium Results (Miami). Tritium
Laboratory Data Rept. #5, Rosenstie1 School of Marine andAtmospheric Science, Univ. of Miami.
All the tritium results obtained on samples from theAtlantic GEOSECS program are given in this report.
52. Ostlund, H.G., H.G. Dorsey, R. Brescher, and W.H. Peterson.
1977. Tritium Laboratory Data Report #6, Oceanic Tritium
Profiles North Atlantic Gyre Sections, 1972-1973, Rosenstiel
School of Marine and Atmospheric Science, Univ. of Miami.
Tritium measurements on profiles taken in 1972 and 1973
along east-west sections at 29°N and 14°N. Also included
in'" this rep 0 rt are the re suI ts from eas tern At Ianti cstations from 14°N to 30 0 S occupied during the same timeinterval. These results compliment those by GEOSECS.
53. Roether, W. and W. Weiss.
section near 40 o N. 1971.

1978.

A transatlantic tritium
Forsch-Ergebn A, V20, p.

"MET~OR"

101-108.

Tritium measurement on samples collected in 1971 along a
traverse across the Atlantic at 40 o N. These results
nicely complement those obtained on samples collectedduring the Atlantic GEOSECS program in 1972.
54. Ostlund, H.G., R. Brescher, R. Oleson and M.J. Ferguson. 1979.

581

GEOSEC.S Pacific radiocarbon and tritium results (Miami).
Tritium Laboratory Data Report #8, Rosenstiel School of Marine
and Atmospheric Science, Univ. of Miami.
All the tritium results obtained on samples from the
Pacific GEOSEGS program are given in this report.

55. Jenkins, W.J., W.V. CDllentro and R.D. Boudreau.

1979. Woods
Hole Oceanographic Institution Helium Isotope Laboratory Data
Release #1, Technical Report Woods Hole Oceanographic InstituInstitution, Woods Hole, Mass.
A listing of tritium and helium 3 results from theSargasso Sea surface waters and thermocline collected overthe period 1974 to 1977.

1980. GEOSECS
Indian Ocean radiocarbon and tritium results. Tritium Laboratory Data Report #9, University of Miami, Miami, Florida.
All the tri tiuro and half of the 14C: resul ts obtained on
samples from the Indian GEOSECS program are given in thisreport.

56. Ostlund, H.G., R. Oleson and R. Brescher.

4. PAPERS DEALING WITH THE INTERPRETATION AND MODELING OF
THE ADDITION OF TRITIUM TO THE SEA AND ITS PENETRATION
INTO THE BODY OF THE SEA (#57 to '70)

57. C~aig, H. and D. Lal. 1961.
tritium. Tellus, V13, p. 85.

The production rate of natural

Estimate of the rate of natural tritium production basedon an integration of the cosmic ray production rates atvarious heights and geographical locations in theatmosphere.

58.

-

W., K.O. Munnich, and H.G. Ostlund. 1970. Tritiumprofile at the North Pacific (1969) GEOSECS intercalibration
station. Jour. of Geophys. Res. V75, p. 7672-7675.
An attempt is made to model the vertical distribution oftritium in the eastern North Pacific using the steadystate advection-diffusion equation.
Roethe~,59. Roath, C.G. and H.G. Ostlund.

1972. Penetration of tritium
into the Atlantic thermocline. Deep Sea Res. V19, p. 481-492.
Concluding from the data in hand that the distribution ofbomb tritium in the north temperate Atlantic thermocline
had reached steady state, the balance between tritium
decay and diffusive mixing is used to obtain upper limitsof the vertical diffusivity in the oceanic thermocline.

60. Fine, R.A. and H.G. Ostlund.

1977. Source function fortritium transport models in the Pacific, Geophys, Res. Lett.

V4,

p.

461-464.

Equations describing the latitude and time histories oftritium in the surface waters of the Pacific Ocean.

61. Dreisigacker, E., and W. Roether.

582

1978.

Tritium and

90 Sr

in

North Atlantic Surface Water. Earth & Planet. Sci. Lett. V38,
p. 301-312.
Two-box model designed to explain the observed temporal
evolution of surface water tritium concentrations in the
North Atlantic Ocean.

62. Hoffert, M.l. and W.S. Broecker.

1978. Apparent verticaleddy diffusion rates in the pycnocline of the Norwegian Sea as
determined from the vertical distribution of tritium. Geophysical Res. Lett. V5, p. 502-504.
A determination of vertical eddy diffusion rates in the
Norwegian Sea based on the vertical distribution oftritium.

63. Weiss, W., W. Roether and E. Dreisigacker. 1979. Tritium in
the North Atlantic Ocean: inventory, input, and transfer into
the deepwater. In: The Behavior of Tritium in the EnvironmentlAEA-SM-232/98.
The primary goal of this paper it to provide estimates ofthe latitude and time dependence of the delivery of bomb
tritium to the surface of the North Atlantic Ocean.
64. Broecker, W.S. and H.G. Ostlund. 1979. Property distributions along the 06=26.8 isopycnal in the Atlantic Ocean.
J. Geophys. Res. V84, p. 1145-1154.
The relationship between the distributions of tritium andof potential temperature on a surface of constant density
in the thermocline of the Atlantic Ocean. The role offronts at 15°N and 15°S is emphasized.
65. Weiss, W. and W. Roether. 1980. The rates of tritium input
to th~ world oceans. Earth & Planet. Sci. Lett. V49, p. 434446.
This paper extends the technique employed by these authors
in reconstructing the tritium input to the North Atlantic
to the other oceans.
66. Fine, R.A. and H.G. Ostlund. 1980. Exchange times in the
Pacific equatorial current system. Earth & Planet. Sci. Lett.
V49, p. 447-452.
A~tempt to estimate the exchange time of water between
the North Equatorial C.ounter C~rrent and the North Equatorial Current based on the distribution of bomb tritium.

67. Quay, P.D., W.S. Broecker, R.H. Hesslein and D.W. Schindler.
1980. Vertical diffusion rates determined by tritium tracer
experiments in the thermocline and hypolimnion of two lakes.
Limnology and Oceanog. V25, p. 201-218.
A discussion of vertical diffusion rates in lakes as
determined by measuring the distribution of purposefultritium tracers and the evolution of the summer temperature profiles with emphasis on the relationship to
density gradient.

583

68. Fine, R.A., J.L. Reid, H.G. Ostlund. 1981. Clrculation of
Tritium in the Pacific Ocean. Jour. of Phys. Ocean., VII, p.
3-14.
Three-dimensional picture of the penetration of tritiumin the thermocline of the Pacific Ocean.
69. Viecelli, J. A., H.W. Ellsaesser and J.E. Burt. 1981. A carbon cycle model with latitude dependence, Climatic Change V3,
p.281-302.
A regional box-diffusion model is constructed using the
distribution of bomb radiocarbon and tritium to calibrate
the mixing parameters.
70. Jenkins, W.J. and P.B. Rhines. 1980. Tritium in the Deep
North Atlantic. Nature, V286, p. 877-880.
Measurement of tritium along the Blake-Bahama Outer Ridge
showing that the southward flowing deep current has
carried bomb tritium to a latitude of 30 o N.

5. PAPERS DEALING WITH THE JOINT INTERPRETATION OF THE
DISTRffiUTIONS OF BOMB RADIOCARBON AND TRITIUM IN THE
OCEAN (171 to '75)

71. Peterson, W.H. and Q.G. Rooth., 1976. Formation and exchange
of deep water in the Greenland and Norwegian Seas. Deep Sea
Res. V23, p. 273-283.
Tritium and radiocarbon data are used to calibrate a modelfor the ventilation of the Greenland and Norwegian Seas.
72. Broecker, W.S., T.-H. Peng and T. Takahashi. 1980. A strategy for the use of bomb-produced radiocarbon as a tracer offossil~ fuel C02 into the deep sea source regions.
Earth &
Planet~ Sci. Lett. V49, p. 463-468.
An attempt is made in this paper to explain the nearlinear relationship between the bomb tritium and bomb
radiocarbon distributions observed for waters in thenorthern Atlantic Ocean.
73. Broecker, W.S. and T.-H. Pengo 1980. The distribution ofbomb-produced tritium and radiocarbon at GEOSECS station 347
in the eastern North Pacific, Earth & Planet. Sci. Lett. V49,
p. 453":'.462.
An attempt to extend the advection-diffusion approach of
Roether et al., to both the vertical distribution ofradiocarbon and tritium at a station in the eastern North
Pacific. As three sets of data exist for this station(Sept. 1969, Nov. 1971 and June 1974), an attempt was madeto match the time distribution as well as the vertical
distributions of these transient tracers.
74. Roether, W., K.O. Munnich, and H. Schoch. 1980. On the 14 C
to tritium relationship in the North Atlantic Ocean. Radiocarbon, V22, p. 636-646.

584

The two-box model previously proposed to explain the temporal history of tritium in the surface waters of theNorth Atlantic is applied to radiocarbon in an attempt to
explain the evolution of the radiocarbon to tritium ratio
in surface waters.

75. Broecker, W.S., T.-H. Peng and R. Engh. 1980.
C~rbon System.
Radiocarbon, V22, p. 565-598.

Modeling the

The application of the distributlton of natural andbomb-produced 14C and of tritium in calibrating ocean
box-diffusion models and regional ocean models.

6. PAPERS DEALING WITH THE DISTRffiUTION OF EXCESS 222 Rn IN
THE DEEP SEA ('76 to #86)

76. Broecker, W.S.

1965. An application of natural radon to problems in ocean circulation. In: Symposium on Diffusion in
Oceans and Fresh Waters, ed. by T. Ichiye, Lamont-Doherty
Geological Observatory, Palisades, New York, p. 116-145.
Discussion of the potential of the radon method for measurement of the vertical mixing rate in near bottom waters.

77. Broecker, W.S., J. Cromwell and Y.-H. Li.

1968. Rates ofvertical eddy diffusion near the ocean floor based on measurements of the distribution of excess 222 Rn . Earth & Planet.
Sci. Lett. V5, p. 101 - 105.
The first paper on near bottom radon measurements and thevertical diffusivities derived therefrom. Also a discussion of the factors controlling the magnitude of the fluxof radon from the sea floor.

78. Broecker, W.S. and A. Kaufman.

1970. Near-surface andnear-bottom radon results for the 1969 North Pacific GEOSECS
Station. Jour. Geophys. Res. V75, p. 7679-7681.
The radon profile at this station shows a distinct maximum
50 meters above the bottom demonstrating the importance ofhorizontal transport from adjacent "hills".

79. Chung, Y.C. and H. Craig.

1972. Excess radon and temperature
profiles from the eastern Equatorial Pacific, Earth & Planet.
Sci. Lett. V14, p. 55-64.
An attempt to relate the distribution of 222 Rn above thesea floor to the pattern of potential temperature change.

80. Sarmiento, J.L., H.W. Feely, W.S. Moore, A.E. Bainbridge and
W.S. Broecker. 1976. The relationship between vertical eddy
diffusion and buoyancy gradient in the deep sea. Earth &Planet. Sci. Lett. V32, p. 357-370.
Demonstration of the inverse relationship between theapparent coefficient of vertical diffusion as establishedfrom the profile of excess 222 Rn and the density gradientin the deep sea .

81. Bainbridge, A.E., P.E. Biscaye, W.S. Broecker, R.M. Horowitz,

585

G. Mathieu, JoL. Sarmiento and D.W. Spencer. 1976. GEOSECS
Atlantic bottom hydrography radon, and suspended particulate
Atlas. GEOSECS Operations Group, Scripps Institution of
Oceanography Internal Report.
Complete listing of deep Atlantic radon and associatedpotential temperature, salinity and density data.
82. Bainbridge, A.E., P.E. Biscaye, R.M. Horowitz, G. Mathieu, J.L. Sarmiento and D.W. Spencer. 1977. GEOSECS Pacific bottomhydrography, radon, and suspended particulate Atlas. GEOSECS
Operations Group, Scripps Institution of Oceanography,
Internal Report.
Complete listing of deep Pacific radon and associatedpotential temperature, salinity and density data.
83. Sarmiento, J.L., W.S. Broecker and P.E. Biscaye. 1978. Excess
bottom radon distribution in deep ocean passages. J. Geophys.
Res. V83, p.5068-5076.
Demonstration that waters passing through oceanic passages
experience very high vertical miXing rates and receive
large amounts of excess radon (presumable from the Mnoxide-lined floor and walls of the passage).
84. Sarmiento, J.L. and C.G.H. Rooth. 1980. A comparison of vertical and isopycnal mixing models in the deep sea based onradon-222 measurements, Jour. Geophys. Res. V85, p. 1515-1518.
An attempt to show that the distribution of excess radon222 in the deep sea might reflect horizontal as well as
vertical mixing processes.
85. Sarmiento, J.L. and W.S. Broecker. 1980. Ocean floor 222 Rn
standing crops in the Atlantic and Pacific Oceans. Earth &Plane~. Sci. Lett. V49, p. 341-350.
Relationship of standing crops of excess 222Rn in thebottom waters of the world ocean to the rate of
sedimentation and to the presence of manganese nodu~es.
86. Kadko, D. 1980. 230 Th , 226 Ra and 222 Rn in abyssal sediments.
Earth & Planet. Sci. Lett. V49, p. 360-380.
An attempt to quantify the relationship between the radon
flux from the sea floor, the rate of sedimentation, theextent of bioturbation and the extent of 226 Ra loss fromthe sediment mixed layer.

7. PAPERS DEALING WITH THE DISTRIBUTION OF 228 Ra IN THE
SEA(187 to #96)

87. Moore, W.S. and W.M. Sackett. 1964. Uranium and thorium series in eqUilibrium in seawater. Jour. of Geophys. Res. V69,
p. 5401-5405.
The discovery of 228 Th and 228 Ra in seawater.
88. Moore, W.S. 1969. Oceanic concentrations of 228Ra.
Planet. Sci. Lett. V6, p. 437-446.

586

Earth &

Preliminary measurements on the distribution of 228 Ra in
the sea.

89. Trier, R.M., W.S. Broecker and H.W. Feely.

1970. Radium-228
profile at the second GEOSECS intercalibration station 1970,
in the North Atlantic. Earth & Planet. Sci. Lett. V16' p.

141-145.

'

Profile of 228 Ra showing surface and bottom maxima.

90. Kaufman, A., R.M. Trier and W.S. Broecker.
of 228 Ra in the world ocean.

8~27-8848.

1973. Distribution
Jour. of Geophys. Res. V78, p.

Data showing the contrast between the 228Ra concentrationin near shore and open ocean surface water samples andestimate of lateral mixing rates derived therefrom.

91. Sarmiento, J.L., H.W. Feely, W.S. Moore, A.E. Bainbridge and
W.S. Broecker. 1976. The relationship between vertical eddy
diffusion and buoyancy gradient in the deep sea. Earth &Planet. Sci. Lett. ¥~2, p. 357-370.
Two deep water
8Ra profiles, one from the North Atlantic
and one from the South Atlantic which yield apparent vertical diffusivity-density gradient products similar to
those obtained from 222 Rn profiles.

1978. Radium andthorium isotopes in the surface water of the east Pacific andcoastal southern CAlifornia. Earth & Planet. Sci. Lett. V39,

92. Knauss, K.G., T.L. Ku and W.S. Moore.
p.

235-240.

Gradients of 228 Ra away from the continental marginand implications regarding rates of lateral mixing.

1979. 228 Th _228 Raradioactive disequilibrium in the New York Bight and its implications for coastal pollution. Earth & Planet. Sci. Lett.

93. Li, Y;-H., H.W. Feely and P.H. Santschi.
V42, p. 13-16.

Gradients of 228Ra away from the continental margin andimplications regarding rates of lateral mixing.

Y.-H., H.W. Feely, and J.R. Toggweiler. 1980. 228Ra and22~Th concentration in GEOSECB Atlantic surface waters. Deep
Sea R~s., V27A, p. ~4~-555.
Distribution of 2 Ra in the upper 350 meters of theAtlantic Ocean.

94. Li95. Moore, W.S., H.W. Feely and Y.-H. Li. 1980. Radium isotopes
in sub-arctic waters. Earth & Planet. Sci. Lett. V49, p.
392-340.
Distribution of 228Ra in the surface waters of theGreenland, Norwegian and Labrador Seas.

96. Sarmiento, J.L., Q.G. Rooth and W.S. Broecker. 1982. Radium228 as a a tracer of isopycnal mixing in the deep ocean Jour.
Geophys. Res.

587

An analysis of the Ra-228 data obtained from the deep
North Atlantic as part of the GEOSECS program suggeststhat isopycnal mixing has a more important influence on
the distribution of this isotope than ·vertical mixing.

8. PAPERS DEALING WITH THE DISTRIBUTION OF 90S~
85 Kr AND 39 Ar IN THE SEA ('97 to 1111)

131 CsJ,

97. Bowen, V.T. and T.T. Sugihara. 1958. Marine geochemical
studies with fallout radioisotopes. In: Proc. Intern. Con.
Peaceful Uses At. Energy, 2nd, Geneva, V18, p. 434.
Measurements of fallout radioisotopes in the Atlantic
Ocean.
98. Bowen, V.T. and T.T. Sugihara. 1960. Strontium-90 in the'mixed layer' of the Atlantic Ocean, Nature, V186, p. 71-72.
Measurements of 90Sr in surface waters from the Atlantic
Ocean.
99. Rocco, G.G. and W.S. Broecker. 1963. The vertical distributionof Cesium-137 and Strontium-90 in the oceans. J. Geophys.
Res., V68, p. 4501-4512.
Measurements of the vertical distribution of 137Cs and
90Sr in the Caribbean Sea showing that as of 1965 theseisotopes had not penetrated beneath 500 meters.
100. Broecker, W.S., E.R. Bonebakker, and G.G. Rocco. 1966. The
vertical distribution of Cesium-137 and Strontium-90 in theocean, 2. J. Geophys. Res. V71, B.1999-2003.
Measurements of 137 Cs and 9 Sr profiles in the North
Atlantic and equatorial Pacific.
101. Bowen,-~V.T., V.E. Noskin, H.L. Volchok and T.T. Sugihara.
1969. Strontium-90: concentrations in surface waters of theAtlantic Ocean. Science V164, p.825-827.
Measurements of 90Sr in samples of Atlantic surface
water.
102. Volchok, H.L., V.T. Bowen, T.R. Folsom, W.S. Broecker, E.A.
Schuert and G.S. Bien. 1971. Oceanic distributions of radionuclides from nuclear explosions. In: Radioactivity in the
Marine~ Environment, edited by A.H. Seymour, U.S. Government
Printing Office, Washington, D.C. p. 42-89.
Summary of the 90 Sr and 137 Cs data for ocean samples up
to 1969.
103. Volchok, H.L. and M.T. Kleinman, 1971, 1972, 1974 and 1975.
Global strontium-90 fallout and precipitation: summary ofdata by 10 degree bands of latitude, Q. Rep. 245, HASL, USAEC
1-2; Q. Rep. 276, HASL, USAEC 1-2; Q. Rep. HASL USAEC 1-17;
Q. Rep. HALS USAEC 1-2.
LatitUde and time dependence of 90Sr fallout as estimatedfrom rainfall and soil measurements.

588

104. Bowen, V.T. and W. Roether. 1973. Vertical distributions ofstrontium-90, cesium-137 and tritium near 45 0 north in the
Atlantic. J. Geophys. Res. V78, p. 6277-6285.
Measurements showing the deep penetration of the fallout
isotopes in the northern Atlantic Ocean.
105. Volchok, H.L. 1974. Is there excess 90 Sr fallout in theoceans? Q. Rep. 286, HASL, USAEC~ 1-82 ..
Addresses the question of the ratio of 90 Sr fallout
over land to that over the ocean.
106. Schroeder, J. 1975. Krypton-B5 in the ocean, Z. Naturforsch,
V30a, p. 962-967.
First report of 85 Kr measurements in sea water.
107. Stockburger, H., H. Sartorius and A. Littkus. 1977.
Messung der Krypton-85 und Xenon-133-Aktivitat der atmospharischen Luft, Z Naturforsch, V32a, p. 1249-125i.
Description of the techniques to measure the ~5Kr and
133 Xe activities in the atmosphere.
108. Dreisigacker, E. and W. Roether. 1978. Tritium and 90 Sr in
North Atlantic surface water. Earth & Planet. Sci. Lett.
V38, p. 301-312.
Relationship between the space and time distributionof tritium and 90 Sr in the North Atlantic.
109. Bowen, V.T., V.E. Noskin, H.D. Livingston and H.L. Volchok.
1980. Fallout radionuclides in the Pacific Ocean: verticaland horizontal distributions, largely from GEOSECS stations.
Earth & Planet. Sci. Lett. V49, p. 411-434.
~Distribution of 137 CB in the North Pacific as determined
Trom samples collected during the GEOSECS program.
110. Weiss, W., A. Sittkus, H. stockburger, H. Sartorius. (inpress) Large-scale atmospheric mixing derived from meridional
profiles of krYPton-B~. Jour. Geophys. Res.
Measurements of 5 Kr in the atmosphere along a south tonorth track in the Atlantic Ocean showing a sharp change
from the ambient southern hemisphere to the higher~mbient northern hemisphere value close to the ITCZ.
Ill. LOOS11, H.H. A dating method with 39 Ar • Earth & Planet.
Sci. Lett. (in press).
First report of 39 Ar measurements in sea water.
B. PAPERS DEALING WITH RADIOTRACERS OF PARTICLE MOVEMENT AND THESCAVENGING OF REAC.TIVE METALS (#112 to #170)

1. PAPERS DEALING WITH THE DISTRIBUTION OF 226 Ra IN THE OCEAN(1112 to #121)
112. Koczy, F.F. 1958. Natural radium as a tracer in the ocean.
Proc. Second U.N. Int. Conf. Peaceful Uses of Atomic Energy589

V18, p. 351-357.
Paper showing how Ra-226 might be used as a tracer ofwater motion.
113. Broecker, W.S., Y.-H. Li and·J. Cromwell. 1967. Radium-226
and radon-222 concentrations in the Atlantic and Pacific
Oceans. Science V158, p. 1307-1310.
Measurements showing the difference in the 226 Ra concentration between waters in the deep Atlantic and in
the deep Pacific.
114. Ku, T.L., Y.-H. Li, G.G. Mathieu and H.K. Wong. 1970. Radiumin the Indian-Antarctic Ocean south of Australia. Jour. ofGeophys. Res. V75, p. 5286-5292.
Distribution of Ra-226 in the Antarctic waters.
115. Chung, Y.-C. 1974. Radium-226 and Ra-Ba relationships in
Antarctic waters. Earth and Planetary Science Lett. V23,
p. 125-135.
Vertical distribution of 226 Ra at several stations in
the Pacific and Antarctic Ocean.
116. Broecker, W.S., J. Goddard and J.L. Sarmiento. 1976. The
distribution of 226 Ra in the Atlantic Ocean. Earth & Planet.
Sci. Lett. V32, p.220-235.
Presentation of the 226 Ra results from the GEOSECS
Atlantic program.
117. Chung, Y.-C. 1976. The deep 226 Ra maximum in the northeast
Pacific. Earth & Planet. ~ci. Lett. V32, p. 249-257.
Presentation of the 2 6 Ra measurements made on GEOSECS~samples from the North Pacific.
118. Chung, Y.-C and H. Craig. 1980. The distribution of radium226 and its relationship with barium and silicate in the
Pacific. Earth & Planet. Sci. Lett. V49, p. 267-292.
Presentation of 226 Ra results from the GEOSECS Pacificprogram.
119. Ku, T.L., C.A. HUb~ and P.S. Chen. 1980. Meridional
distributions of 2L6 Ra in the eastern Pacific along GEOSECS
cruise tracks. Earth & Planet. Sci. Lett. V49, p. 293-308D
Presentation of 226 Ra results from the GEOSECS Pacificprogram.
120. Kadko, D. 1980. 230 Th , 226 Ra , and 222 Rn in abyssal sediments. Earth & Planet. Sci. Lett. V49, p. 360-380.
A discussion of 222 Rn flux out of deep sea sediments.
This flux is related to the 226 Ra profile in thesediment which in turn depends on the 230 Th rain rate
and the bulk sedimentation rate.
121. Cochran, J.K. 1980. The flux of 226 Ra from deep sea sediments. Earth & Planet. Sci. Lett. V49, p. 381-382.

590

Estimates of the flux of 226Ra from deep sea sediments
based on 226Ra/230 Th ratio measurements. The results
show far higher fluxes for cores from the eastern Pacificthan for those from the Atlantic.
2. PAPERS DEALING WITH THE DISTRffiUTION OF 234 Th ,230Th

AND 231pa IN THE OCEAN (1122 to 1138)

228Th

,

122. Moore, W.S. and W.M. Sackett.

1964. Uranium and thorium
series inequilibrium in sea water. Jour. Geophys. Res. V69,
p. 5401-5405.
The demonstration of the lack of equilibrium between
228 Th and 228 Ra •123. Kaufman, A. 1969.

The 232Th concentration of surface ocean
water. Geochim. Cosmochim. Acta V33, p. I17-724.
Upper limit on the concentration of 2 2Th in surface
ocean water.

124. Bhat, S.G., S. Krishnaswami, D. Lal, M. Rama, and W.S. Moore.
1969. 234 Th /238 ratios in the ocean. Earth and Planetary
Science Lett., V5, p. 483-491.
Demonstration of the lack of equilibrium between 234 Th
and 238U in surface waters.

125. Broecker, W.S., A. Kaufman, and R.M. Trier.

1973. The residence time of tho~ium in surface seawater and its implications regarding the fate of reactive pollutants~ Earth andPlanet. Sci. Lett. V20, p. 35-44.
228Th to 228 Ra measurements in surface ocean water.

126. Matsumoto, E. 1975.
layer of the ocean.

212. -

234Thj238U disequilibrium in the surface
Geochim. Cosmochim. Acta, V39, p. 205-

Measurements showing the deficiency of 234 Th in surface
waters.

127. Ku, T.L., K.G. Knauss, and G.G. Mathieu.

1977. Uranium inopen ocean: concentration and isotopic compensation. Deep
Sea Res. V24, p. 1005-1007.
Summary of data demonstrating the constancy of the uranium content of sea salt and of the 234Uj238U in sea water.

128.

Spence~, D.W., P.G. Brewer, A.P. Fleer, S. Honjo, S. Krishnaswami and Y. Nozaki. 1978. Chemical fluxes from a sediment trap experiment in the deep Sargasso Sea. J. Mar. Res.
V36, p. 493-523.
Thorium isotope concentrations in the material caught insediment traps and their implications to trapping efficiency and particulate scavenging rates.

129. DeMaster, D.J.

1979. The marine bUdgets of silica and328i . Ph.D. thesis, Yale University, New Haven, Conn.
Measurements of 231pa and 230Th concentrations in sediment core tops showing that 231 pa is preferentially591

deposited beneath areas of upwelling.

130. Li, Y.-H., H.W. Feely and P.H. Santschi.

1979. 228 Th -228 Raradioactive disequilibrium in the New York Bight and its implications for coastal pollution. Earth and Planet. Sci.
Lett., V42, p. 13-26.
Measurements showing the decrease in 228 Th residence
time in surface waters as the coast is approached.

131. Minagawa, M. and S. Tsunogai.

1980. Removal of 234 Th from acoastal sea: Funka Bay. Earth and Planet. Sci. Lett., V47,
p. 51-64.
Measurements demonstrating the rapid removal of thoriumfrom coastal waters.

132.

Li~ Y.-H., H.W. Feely and J.R. Toggweiler.
1980. 228 Ra and22~Th concentrations in GEOSECB Atlantic surface waters. Deep

Sea Research, V27A, p. 545-5SS.
.
Demonstration that the 118 Th extracted from surface
waters must be in part returned to solution in the upper
thermocline.

133. Santschi, P.H., D. Adler, M. Amdurer, Y.-H. Li and J. Bell.
1980. Thorium isotopes as analogues for "particle-reactive"
pollutants in coastal marine environments. Earth andPlanet. Sci. Lett. V47, p. 327-335.
C.omparison between the removal rates of thorium isotopes
in Narragansett Bay and in controlled ecosystems designed to duplicate the conditions in this water body.

134. Anderson, R.F.

1981. The marine geochemistry of thorium andprotactinium. Ph.D. thesis. Woods Hole Oceanographic Institution, Woods Hole, Mass.
Measurements of thorium and protactinium isotopes inmaterials from sediment traps, filtered particulates,
and manganese adsorbers at four ocean stations.

135. Nozaki, Y., Y. Horibe and H. Tsubota.

1981. The watercolumn distributions of thorium isotopes in the western North
Pacific. Earth and Planet. Sci. Lett. V54, p. 203-216.
~easurements of 230Th dissolved in sea water as a function of depth in the North Pacific Ocean. A first!
Also a demonstration that the adsorption of thorium onto
particles must be reversible.

136. Moore, W.S.

1981. The thorium isotope content of ocean
water. Earth & Planet. Sci. Lett. V53, p. 419-426.
Measurements of the 230 Th content of surface and deep
water showing that the residence time increases withdepth reaching as much as 100 years in near bottom
waters.

137. Krishnaswami, S., M.M. Sarin and B.L.K. Somayajulu.

1981.

Chemical and radiochemical investigations of surface and deep592

particles in the Indian Ocean. Earth & Planet. Sci. Lett.
V54, p. 81-96.
'
Measurements of the vertical distribution of the amountof 230 Th on suspended particulate matter in the Indian
Ocean.
138. Kaufman, A' Y.-H. Li and K.K. Turekian. 1981. The removal
rates of 23 4Th and 228 Th from waters of the New York Bight.
Earth & Planet. Sci. Lett., V54, p. 385-392.
Comparison of thorium removal times calculated from
234 Th /238 U and 228 Th /228 Ra ratios.
138a.Bacon, M.P. and R.F. Anderson. 1982. Distribution ofthorium isotopes between dissolved and particulate forms inthe deep sea. Jour. Geophys. Res. V87, p.2045-2056.
Demonstration that the fraction of thorium on particulates depend on the half-life of the isotope and on theconcentration of the suspended matter.

3. PAPERS DEALING WITH THE DISTRffi UTION OF 2 10 Po AND 21 °pb
IN THE SEA (1139-1161)
139. Rama, M., Koide, M. and E.D. Goldberg. 1961. Lead-210 innatural waters. Science V134 p. 98-99.
First paper showing that 2tO Pb from atmospheric fallout ispresent in natural waters.
140. Fisenne, I.M. 1968. Distributions of lead-210 and radium226 in soil. Rep. UCRL-18140 U.S. Dept. of Energy, Wash.,
D.C:. p. 145-158.
Use of soil data to estimate the flux of 210 Pb atoms outof the atmosphere.
140a.Windom, H.L. 1969. Atmospheric dust records in permanentsnowfields: Implications to marine sedimentation. Geol. Soc.
Amer. Bull. V80, p.761-782.
Evaluation of the contribution of atmospherically transported solids to recent marine sediments based on the dust
accumulation in the permanent snowfields.
141. Shannon, L.V., R.D. Cherry and M.J. Orren. 1970. Polonium210 and lead-210 in the marine environment. Geochim. Cosmochim. 'Acta, V34, p. 98-99.
210 Pb concentrations in marine plankton.
142. Tsunogai, S. and Y. Nozaki. 1971. Lead-210 and polonium-210
in the surface water of the Pacific. Geochem. J., V5, p. 165.
Distributions of 210 pb and 210 po in the sea.
142a.Krishnaswami, S., D. Lal, J.M. Martin, M. Meybeck. 1971.
Geochronology of lake sediments. Earth & Planet. Sci. Letts.
VII, p. 407-414.
This paper describes partially successful attem~gs tomeasure rates of sedimentation in lakes using 2 Pb

593

Y19hnique. Other radiometric techniques including
Cs and 32Si are also described.

55 Fe ,

143. Nozaki, Y. and S. Tsunogai. 1973. Lead-210 in the North
Pacific and the transport of terrestrial material through theatmosphere. Earth & Planet. Sci. Lett. V20, p. 88-92.
Distribution of 210 Pb in the North Pacific Ocean.
144. Beasly, T.M., R.J. Eagle and T.A. Jokela. 1973. Polonium210, lead-210 and stable lead in marine organisms. U.S.
Atom~1oEnergy Commission, HASL-273, p I .
Pb concentrations in marine plankton.
145. ~1~ig'2~6' S. Krishanaswami and B.L.K. Somayajulu. 1973.
PbRa: radioactive disequilibrium in the deep sea.
Earth & Planet. Sci. Lett. ¥17, p. 2~5-305.
Determination of the 21 Pb to 22 Ra activity ratio in
sea water.
145a.Gavini, M.B., J.N. Beck, and P.K. Kuroda. 1974. Mean residence time of the long-lived radon daughters in the atmosphere. J. Geophys. Res. V79, p. 4447-44~2.
Measurements of the activities of 21 Pb, 210 Bi and 210 poin rain samples at Fayetteville, Arkansas. Mean residence time of these nuclides in the atmosphere are calculated.
146. Bruland, K.W., M. Koide and E.D. Goldberg. 1974. The comparative marine geochemistries of lead-210 and radium-226.
J. Geophys. Res. V79~~.3083-3086.
Measurements of
6 Ra and 210 Pb concentration in the Gulf
of California and eastern North Pacific waters.
146a.Robbins, J.A. and D.N. Edgington. 1975. Determina~1~n ofrecent sedimentation rates in Lake Michigan using
Pb and137 Cs .
Geochirn. Cosmochim. Acta V39 J ~. 285-30i.
This paper describes the use of 21 Pb and 1 7 Cs radioactivity measurements to determine the rates of sedimentation in large lakes.
147.

S., B.L.K. Somayajulu and Y.-C. Chung. 1975.
210 Pb /,.2 2 Ra disequilibrium in the Santa Barbara Basin. Earth& Planet. Sci. Lett. V27, p. 388-392.
Measurements of 210 Pb and 226 Ra in a small coastal basin.

Krishnas~ami,147a.Kipphut, G.W. 1978. An investigation of sedimentary processes in lakes. Ph.D. thesis, Columbia Dniv., LamontDoherty Geological Obs., Palisades, N.Y.
A discussion of the rates of sediment accumulation, rates
and extents of sediment mixing, and rates of trace metal
accumulation in the sediments of small lakes.
147b.Matsumoto, E. 1973. 210 Pb geochronology of sediments from
Lake Shinji. Geochem. Jour. V9, p. 167-172.

594

210
ExcessPb is used successfully for dating lake sediment cores.
148. Soma~a~ulu, B.L.K. and H. Qraig. 1976. Particulate and soluble 1 Pb activities in the deep sea. Earth & Planet. Sci.
Lett. V32, p. 268-276.
Comparison of the amount of 210 Pb on particles andthe amount dissolved in sea water.
149. Bacon, M.P., D.W. Spencer and P.G. Brewer. 1976.
210 Pb /226 Ra and 210 po/210 Pb disequilibria in sea water andsuspended particulate matter. Earth & Planet. Sci. Lett.
V32, p. 277-296.
Measurements of 210 po , 210 Pb and 226 Ra in seawater and insuspended particulates.
150. Thomson, J. and K.K. Turekian. 1976. 210 po and 210 pb distributions in ocean water from the eastern South Pacific.

Eart~h: ~i:~~;bU~~~~S~;tilo~~2~ng·2f~~~3~~·the sea.
151. Nozaki, Y., J. Thomson and K.K. Turekian. 1976. The distributions of 210Pb and 210po in the surface waters of the
Pacific Ocean. Earth & Planet. Sci. Lett. V32, p. 304-312.
The distributions of 210 po and 210 Pb in the sea.
152. Nozaki, Y. and S. Tsunogai. 1976. 226 Ra , 210 Pb , and 210 po
disequilibria in the western North Pacific. Earth & Planet.
Sci. Lett., V32, p. 313-321.
The distributions of 226 Ra , 210 Pb , and 210 po in the sea.
153. Moore~ H.E. and S.E. Poet. 1976. 210 Pb fluxes determinedfrom -,loPb and 226 Ra soil profiles. J. Geophys. Res . V81,
p. 1056-1058.
Use of soil data to estimate the flux of 210Pb atoms
from the atmosphere.
154. Turekian, K.K., Y. Nozaki and L.K. Benninger. 1977. Geochemistry of atmospheric radon and radon products. Ann.
Rev. Earth & Planet. Sci. V5, p. 227-255.
~eview paper discussin~ the sources of atmospheric radon
and the sinks for the lOpb produced by the decay ofradon in the atmosphere.
155. Nozaki, Yo, D.J. DeMaster, D.M. Lewes and K.K. Turekian.
1978. Atmospheric 210 Pb fluxes determined from solI profiles. Jour. of Geophys. Res. V83, p. 4047-40~1.
Use of soil data to estimate the flux of 210 pb atoms fromthe atmosphere.
155a.Spencer, D.W., P.G. Brewer, A. Fleer, S. Honjo, S. Krishnaswami and Y. Nozaki. 1978. Chemical fluxes from a sediment
trap experiment in the deep Sargasso Sea. Jour. Mar. Res.
V36, p. 493-523.

595

Measurements of the flux of 210 Pb atoms toward the deepsea floor.
156. Beasly, T.M., M21~eyraud'21oJ.W. Higgo, R.D. Cherry and S.W.
Fowler. 1978.
Po andPb in zooplankton fecal pellets. Mar. BioI. V44, p. 325-328.
210 Pb in waste products of organisms.
157. Tsunogai,S. and K. Harada. 1980. 226 Ra and 210 Pb in thewestern North Pacific. In: Isotope Marine Chemistry, ed. by
E.D. Goldberg, Y. Horibe and K. Saruhashi, Uchida Rokahu Ho,Tokyo, p. 165-191.
Profiles of 210Pb and 226 Ra in the North Pacific Ocean.
158. Nozaki, Y., K.K. Turekian and K. von Damm. 1980. 210 Pb in
GEOSECS water profiles from the North Pacific, Earth &Plan~~~ ~~;trt~~~io~4;f ~io~~3i~0~he sea.
159. Turekian, K.K. and Y. Nozaki. 1980. 210 po and 210 Pb in theeastern South Pacific. The role of upwelling on their distribution in the water column. In: Isotope Marine Chemistry,ed. by E.D. Goldberg, Y. Horibe and K. Saruhashi, Uchida
Rokahu Ho, Tokyo, p. 157-164.
The distribution of 210 Pb in the sea.
160. Chung, Y. and M.D. Applequist. 19BO. 226 Ra and 210 pb in the
Weddell Sea. Earth & Planet. Sci. Lett. V49, p. 401-410.
The distribution of 210 pb in the Antarctic Ocean.
161. Moore, W.S., K.W. Bruland, and J. Michel. 1981. Fluxes ofuranium and thorium series isotopes in the Santa Barbara
Basin~
Earth & Planet. Sci. Lett. V53, p. 391-399.
Demonstration that the 210 Pb flux to the sediments of thesanta Barbara Basin is about seven times the sum of thewater column production and atmospheric fallout rate.
210 Pb from the adjacent surface ocean must enter the
Santa Barbara basin and be removed onto particles there.
161a.Spencer,

D.W.~

M.P. Bacon and P.G. Brewer. 1980. The dislOPb and 210 po in the North Sea. Thalassia
Jugoslavia V16, p. 125-154.
Strong depletions in 210Pb content are found in North Sea
waters.
trib~tion of

4. PAPERS DEALING WITH THE INPUT OF AND DISTRIBUTION OF
PLUTONIUM WITHIN THE SEA (#162 TO 1165)

162. Hardy, E.P., P.W. Krey and H.L. Volchok. 1973. Global
inventory and distribution of fallout plutonium. Nature,
V24l, p. 444-445.
Estimate of the amount and latitude dependence of
plutonium fallout.

596163. Noskln, V.E., and K.M. Wong. 1980. Plutonium in the North
Equatorial Pacific. In: Processes Determining the Input Behaviour and Fate of Radionuclides and Trace Elements inContinental Shelf Environments, U.S. Dept. of Energy, Rep.
C.ONF-790382, p. 11-28.
Measurments of plutonium in sea water samples from the
Pacific.
164. Bowen, V.T., V.E. Noskin, H.D. Livingston and H.L. Vo1chok.
1980. Fallout Radionuclides in the Pacific Ocean: verticaland horizontal distributions, largely from GEOSECB stations.
Earth & Planet. Sci. Lett. V49, p. 411-434.
The vertical distribution of 239pu and 2qOpo in the
North Pacific. Evidence for regeneration within thethermocline.
165. Santschi, P.H. and Y.-H. Li. Removal Pathways of Th and Pu
Isotopes in C~astal Marine Environments. Proceedings of theSecond Symposium on Natural Radiation Environment, Bombay,
India. D.O.E. Symposium Series, (in press).
The correlation between plutonium and thorium in the
coastal waters off New York suggest that these twoelements are removed by the same mechanism. Evidence
is presented that 228 Th is removed by adsorption onto
particles.

5. PAPERS DEALING WITH THE DISTRIBUTIONS OF 14 C IN PARTI~ULATE DEBRIS AND ORGANISMS (#166 to #167)
166. Somazajulu, B.L.K., D. Lal, and S. Kusumger. 1966. Man-madecarbon-14 in deep Pacific waters: transport by biologicalskeletal material. Science, V166, p. 1379-1399.
Demonstration that the CaC:03 particles falling throughthe deep sea areof quite recent origin based on thepresence of bomb-produced radiocarbon in samples ofsuspended particles filtered from deep sea water.
167.

Williams, P.M., J.A. McGowan and M. Stuiver. 1970. Bombcarbon-14 in deep sea organisms. Nature, V227, p. 375-376.
Demonstration that the food eaten by deep dwelling fish
is of recent origin through the finding that the flesh oft~ese organisms contains bomb-produced radio carbon.

6. MODELS FOR PARTICnLATE SCAVENGING IN THE OCEAN (#168
to #170)

16e. Craig, H. 1974. A scavenging model for trace elements in the
deep sea. Earth & Planet. Sci. Lett. V23, P. 149-159.
Application of the one dimensional advection-diffusion
model to trace metals. A scavenging term which assumes
first order irreversible removal onto particlels is added
to the water transport equations. Unlike the nutrient597

elements which show excesses when plotted against
potential temperature, the metals copper, antimony andscandium show deficiencies suggesting that they are being
scavenged from deep water faster than they are being
added by particle destruction.

169. Tsunogai, S. and M. Minagawa.

1978. Settling model for theremoval of insoluble chemical elements in sea water. Geochern. J., V12, p. 47-56.
Modeling the removal mechanism of chemical elements inthe ocean based on the behavior of insoluble radionuclides in seawater.

170. Spencer, D. W., M.P. Bacon, and P .G. Brewer.

1981. Models of
the distribution of 210 Pb in a section across the north equatorial Atlantic Ocean. Jour. of Mar. Res. V~~t p. 119-138.
Two-dimensional model of the removal ofPb from theocean involving enhanced removal near the boundaries.
This enhanced removal could either be the result of direct uptake by the sediment or the result of higher particulate content of the water near the oceans' boundaries.

C. PAPERS DEALING WITH RADIO TRACERS OF GAS EXClIANGE (#171 to'194)
1. PAPERS DEALING WITH THE RADON METHOD

171. Broecker, W.S.

(1171 to 1177)

1965.

An applcation of natural radon to
problems in ocean circulation. In: Symposium on Diffusion in
Oceans and Fresh Waters. Lamont-Doherty Geological Observatory, Palisades, NoY., p. 116-145.
First publication showing the potential of radon as a
gas-exchange tracer.

172. Broecker W.S. and T.-H. Pengo

ion of radon in the BOMEX area.

1971.

The vertical distribuEarth & Planet. Sci. Lett.

VII, p. 99-108.

The first field application of the radon method. A longtime series of 222 Rn and 226 Ra measurements in the mixed
layer and upper thermocline in the open Atlantic off theisland of Barbados permitted the establishment of the gas
e~change rate under trade-wind conditions.

173. Peng, T.-H., T. Takahashi and W.S. Broecker.

1974. Surfaceradon measurements in the North Pacific Ocean station Papa.
J. Geophys. Res. V79, D. 1772-1780.
A time series of 222 Rn and 22bRa measurements understormy winter conditions at 50 0 N in the Pacific Ocean.

174. Emerson, S. 1975.

Gas exchange rates in small Canadianshield lakes. Limnol. Oceanogr., V20, p. 754-761.
Addition of a 226 Ra s~ike to the mixed layer of smalllake to enhance the 2 2Rn production and thereby allow
estimates of the gas transfer rate to be made.

598

175. Bainbridge, A.E., W.S. Broecker, R.M. Horowitz, Y.-H. Li, G.
Mathieu and J. Sarmiento. 1977. GEOSECS Atlantic andPacific Surface Hydrography and Radon Atlas. GOG Publication#120, Scripps Institution of Oceanography, La Jolla, Calif.
Report containing listings of the radon and associatedhydrographic data for the Atlantic and Pacific GEOSECS
expeditions.
176. Peng, T.-H., W.S. Broecker, G.G. Mathieu and Y.-H. Li. 1979.
Radon evasion rates in the Atlantic and Pacific Oceans as determined during the GEOSECS program. Jour. of Geophys. Res.,
Ve4, ~. 2471-24~~.
2 2Rn and 2 Ra measurements for surface mixed layer andupper thermocline samples from the Atlantic, Pacific andAntarctic Oceans. The gas exchange velocities obtainedin this way permit regional and global averages to beobtained.
177. Kromer, B. and W. Roether. 1980. Wind dependence of gas exchange determined in field experiments with the radon method.
Symposium on "Capillary waves and gas exchange", Trier 1979,
pUblished by SFB94 Meeresforschung, Hamburg, p. 87-90.
Use of integrated 222 Rn profiles to estimate the rate ofgas exchange.
2. PAPERS DEALING WITH THE 3H/3 He METHOD (1178 to #179)178. ~orgersen, T., Z. Top, W.B. C~arke, W.J. Jenkins, W.S.
Broecker. 1977. A new method for physical limnology;
tritium/helium-3 ages; results for Lakes Erie, Huron, andOntario. Limnol. Oceanog. V22, p. 181-193.
»emonstration that the helium-3 to tritium ratio in themixed layer of fresh water lakes can be used to determine the gas exchange rate.
179. Torgerson, T., G. Mathieu, R.H. Hesslein and W.S. Broecker.
1982. Gas exchange dependency on diffusion coefficient:
Direct 222 Rn and 3 He comparisons in a small lake. J. Geophys. Res. V87, p. 546-556.
The mixed layer of a small Canadian shield lake was
spiked with tritium and with 226 Ra permittin~ the simultaneous measurement of the evasion fluxes of He and
2.22Rn. Under these low-wind conditions the stagnant
film model appears to provide a better approximation to
the gas exchange process than the film replacement model.
3. PAPERS DEALING WITH THE RADIOCARBON METHOD (#180 to1184)

180. Craig, H. 1957. The natural distribution of radiocarbon andthe exchange time of carbon dioxide between the atmosphereand the sea. Tellus, V9, p. 1-17.
Method of determining the average C02 exchange rate between air and sea based on the distribution of naturalradiocarbon.

181. Munnich, K.O. and W. Roether. 1967. Transfer of bomb 14 C
and tritium from the atmosphere to the ocean. Internal mixing of the ocean on the basis of tritium and 14 C profiles.
In: Radioactive Dating and Methods of Low Level Counting.
Proceedings IAEA-ICSU Symposium, Vienna, V92, p. 93-104.
Use of the ocean inventory of bomb-produced radiocarbon
to determine the CO 2 invasion rate.
182. Broecker, W.S., T.-H. Peng and R. Engh. 1980. Modeling of
the carbon system. Radiocarbon, V22, p. 565-598.
Use of the inventory of bomb 14 C in the ocean to estimate the rate of CO 2 invasion.

183. Hesslein, R., W.S. Broecker, P.D. Quay and D.W. Schindler.
1980. Whole lake radiocarbon experiment in an oligotrophic
lake at the Experimental Lakes Area Northwestern Ontario.
Canadian J. Fish. Aquatic Sci. V37, p. 454-463.
A discussion of carbon cycling in lakes based on dataobtained by spiking the whole lake with radiocarbon.
184. Peng, T.-H. and W.S. Broecker. 1980. Gas exchange rates forthree closed-basin lakes. Limnol. and Oceanog. V25, p. 789796.
Utilization of the bomb 14C~ inventory in saline lakes
from Great Basin of the western U.S. to determine theC02invasion rates.

4. PAPERS DEALING WITH LABORATORY MEASUREMENTS OF GASEXCHANGE RATES (N185 to 1188)
185. Downing, A.L. and G.A. Truesdale.

1955. Some factors afthe rate of solution of oxygen in water. J. Appl.
Chern.; V5, p. 570-581.
Demonstration of the role of wind velocity in determining the rate of gas exchange.
fect~ng

Kanwisher, J. 1963. On the exchange of gases between theatmosphere and the sea. Deep Sea Res., VI0, p. 195-207.
Lab experiments showing the effect of wind velocity andwaves on the exchange rate of CD2 between air and sea
w.ater.
187. Liss, P.S. 1973. Processes of gas exchange across an airwater interface. Deep Sea Res. V20, p. 221-238.
A review of our knowledge of the factors controlling
the rate of gas exchange between the atmosphere andnatural waters.
188. Broecker, H.-C~., J. Peterman and W. Siems. 1979. The influence of wind on C02-exchange in a wind-wave tunnel,
including the effects of monolayers. Jour. of Marine Res.,
V36, p. 595-610.
Wind tunnel measurements demonstrating the importanceof capillary waves on the rate of gas exchange.

600

5. THEORETICAL TREATMENTS OF VARIOUS ASPEC~S OF THE GASEXCHANGE PROCESS (1189 to 1194)
189. Bolin, B. 1960. On the exchange of carbon dioxide between
the atmosphere and the sea. Tellus, V12, p. 274-281.
Discussion of the role of the reactions between C02 andH20 and CO 2 and OH- in the enhancement of CO 2 exchange.
A classic paper in the field of CO 2 exchange.
190. Danckwerts, P.V. 1970. Gas-liquid reactions. MCGraw-Hill,
New York.
The definitive chemical engineering text on gas
exchange theory and experiment.
191. Munnich, K.O. and D. Flothmann. 1975. Gas exchange in relation to other air/sea interaction phenomena. SCDR Workshop on"Air/Sea Interaction Phenomena" Miami.
The relationship between momentum, heat and gas
transfer.
192. Emerson S. 1975. Chemically enhanced C02 gas exchange in aeutrophic lake: a general model. Limnol. and Oceanog. V20,
p. 743-753.
Demonstration that under the high pH conditions generated by excess algal growth in dissolved carbon-poorlakes a significant enhancement (over that for othergases) in the rate of C02 invasion occurs.
193. Deacon, E.L. 1977. Gas transger to and across an airWater interface. Tellus V29, p. 363-374.
Theoretical treatment of the gas exchange process sug_gesting that the film replacement model should provide aoetter approximation than the stagnant film model.
194. Munnich, K.O., W.B. Clarke, K.H. Fisher, D. Flothmann, B.
Kromer, W. Roether, U. Siegenthaler, Z. Top and W. Weiss.
1978. Gas exchange and evaporation studies in a circular
wind tunnel, continuous radon-222 measurements at sea andtritium/helium-3 measurements in a lake. In: Turbulent
fluxes through the sea surface, wave dynamics and prediction,edited by A. Favre and K. Hasselmann, Plenum PUblishing~orpdration, p. 151-166.
By measuring both the rate of heat transfer and gas
transfer in a circular wind it was shown that thefilm replacement model yields a better approximation
than the stagnant film model.

D. PAPERS DEALING WITH RADIOISOTOPES IN OCEAN SEDIMENTS (#195
to #264)

1. PAPERS DEALING WITH RADIOCARBON MEASUREMENTS IN DEEPSEA SEDIMENTS (D195 to #199)
195. Broecker, W.S., K.K. Turekian and B.C. Heezen. 1958. The

601

relation of deep sea sedimentation rates to variations inclimate. American Jour. Sci., V256, p. 503-5170
Detailed dating of a deep sea core from the equatorial
Atlantic showing a large change in the rate of deposition
between glacial and interglacial time for both the clay
and carbonate fractions.

196. Ericson, D.B., M. Ewing, G. Wollin, and B.C. Heezen. 1961.
Atlantic deep sea sediment cores. Bull. Geol. Soc. Am., V72,
p.193-285.
Radiocarbon dates for the end of the last glaciation asrecorded in deep sea cores.

197. Peng, T.-H., W.S. Broecker, G. Kipphut and N. Shackleton.
1977. Benthic mixing in deep sea cores as determined by 14 C
dating and its implications regarding climate stratigraphy
and the fate of fossil fuel CO 2 - In: "The Fate of Fossil
Fuel CD2 in the Ocean", editors N. Andersen and A. Malahoff,
Plenum Press, New York, p. 355-374.
One of the two independent discoveries that the cacol in
the upper 7-10 em of deep sea cores yields constant 4 C
ages of several thousands of years and that this anomaly
must be the result of bioturbation.

198. Nozaki, Y., J.K. cochran 1 K.K. Turekian and G. Keller.
1977. Radiocarbon and 2 °Pb distributions in submersibletaken deep sea cores from Project FAMOUS. Earth & Planet.
Sci. Lett. V34, p. 167-173.
One of the two independent discoveries that the C~~03 in
the upper 4-10 centimeters of deep sea cores yields
constant 1 C ages of several thousand years and that thisanomaly must be the result of bioturbation.

1979. Rates ofbenthic mixing in deep-sea sediment as determined by radioactive tracers. Quaternary Res. VII, p. 141-149.
Another example of the use of radiocarbon measurements
to establish the thickness of the bioturbated zone indeep sea cores.

199. Peng, T.-H., W.S. Broecker and W.H. Berger.

~.

PAPERS DEALING WITH THE 230 Th AND 231 pa DATA IN DEEPSEA SEDIMENTS AND MANGANESE NODULES (#200 to #239)

200. Joly, J.J. 1908. On the radium content of deep sea sediments. Philos. Mag. V6, p. 190-197.
The earliest report of measurements of uranium series
daughter products in marine sediments.
201. Piggot, C.. S. 1933. Radium content of ocean bottom sediments. Am. J. Sci. V239, ~. 81-91.
First measurement of 26 Ra in a deep sea sediment.

202. Piggot, C.. S. and W.D. Urry.
ocean bottom core.

1939.

The radium content of an

J. Wash. Acad. Sci. V29, p. 405-410.

602

Use of the depth distribution of 226 Ra in deep sea coresas a means of estimating the 230 Th depth distribution
and hence the sedimentation rate by the ionium method.
203. Piggot, C.S., and W.D. Urry. 1941. Radioactivity relationsin ocean water and bottom sediments. Am. J. Sci. V239, p.
81-91.
The geochemistry of 226 Ra and 230 Th in the sea.
204. Piggot, C.S. and W.D. Urry. 1942. Time relations in ocean
sediments. Bull. Geol. Soc. Amer. V53, p. 1187-1210.
Age estimates based on the 226 Ra distribution in deepsea sediments.
205. Piggot, C.S. and W.D. Urry. 1942. Radioactivity of ocean
sediments IV. The radium content of sediments of the Cayman
Trough. Am. J. Sci. V240, p. 1-12.
226 Ra distribution in cores from the Caribbean Sea.
206. Kroll, V.S. 1953. Vertical distribution of radium indeep sea sediments. Nature, V171, p. 742.
226 Ra distribution in deep sea cores.
207. Kroll, V.S. 1954. On the age-determination in deep sea
sediments by radium measurements. Deep-Sea Res. VI, p.
211-215.
Ages obtained from the 226 Ra distribution in deep sea
cores.
208. Picciotto, E. and W. Wilgain. 1954. Thorium deteNninationin deep sea sediments. Nature V17i~ p. 632-633.
~ First direct measurement of 2juTh in deep sea sediments.
209. Koczy, F.F., E. Picciotto, G. Poulart and S. Wilgain. 1957.
Measures des isotopes du thorium dans l'eau de mer. Geochim.
Cosmochim. Acta, VII p. 103-12~.
Direct measurements of 23 Th in deep-sea cores.
210. Volchok, H.L. and J.L. Kulp. 1957. The ionium method of age
determination. Geochim. et Cosmochim. Acta, VII, p. 219-246.
More 226 Ra measurements on deep-sea cores.
211. Emi11ani, C. 1958. Paleotemperature analysis of core 280 andPleistocene correlations, J. Geol. V66, p. 264-275.
Correlations of paleotemperature record in core 280 frommid-North Atlantic with cores from equatorial Atlantic,
Caribbean, and Mediterranean, and with continental
temperature variations during last 300,000 years.
212. Sackett, W.M. 1960. The protactinium-231 content of ocean
water and sediments. Science V132, p. 1761-1762.
First measurement of 231 pa in marine sediments.
213. Rosholt, J.N., C. Emiliani, T. Geiss, F.F. Koczyand

603

J. Wan~ersky. 1961. Absolute dating of deep-sea cores by the
231 pa / 30Th method. J. Geol. V69, p. 162-18S.
First simultaneous measurements of 23~Th and 231 pa indeep sea cores and first ages based on 230 Th /231 pa
ratios. These ages were used to establish what isoften referred to as the Emiliani chronology for thelast several major climatic cycles.
214. Goldberg, E.D. and M. Koide. 1962. Geochronological studies
of deep sea sediments by the ionium/thorium method. Geochim.
C~osmochim. Acta, V26 p. 417-450.
Sedimentation rates determined using 230 Th /232 Th
ratios on material leached from deep sea sedimentswith strong acid.
215. Goldberg, E.D. and M. Koide. 1963. Rates of sediment accumulation in the Indian Ocean. In: Earth Sci. & Meteorites,
Ed. J. Geiss, North Holland, Amsterdam, p. 90-95.
230Th based sedimentation rates for cores from theIndian Ocean.
216. Thurber, D.L. 1963. Natural variations in the ratio of
234U and 238U; In: "Radioactive Dating," Internat. Atomic
Energy Agency, Vienna, p. 113-120.
First demonstration of the existence of a 234U/238U
in sea water.
Sarma, T.P. 1964. Dating of marine sediments by ionium andprotactinium methods. Ph.D. Thesis, Carnegie Institute ofTechnology.
First suggestion that bioturbation altered the
profiles of uranium series isotopes in deep sea cores.
218. Goldberg, E.D., M. Koide, J.J. Griffin, M.N.A. Peterson.
1964. A geochronological and sedimentary profile across theNorth Atlantic Ocean. In: Isotopic and Qosmic Chemistry,ed. by H. Craig, S.L. Miller and G.J. Wasserberg, NorthHolland~ Amsterdam, p. 211-232.
2 0Th profiles in cores showing large changes in
sedimentation rate.
Goldberg, E.D. and J.J. Griffin. 1964. Sedimentation rates
and mineralogy in the South Atlantic. J. Geophys. Res. V69,
p. 4293-4309.
Sedimentation rates on South Atlantic cores as
determined by the 230 Th /232 Th method.
220. Sackett, W.M. 1966. Manganese nodules: thorium-230:
protactinium-231 ratios. Science, Vl~4, p. 646-647.
Demonstration that the 231 pa /2 °Th ratio is higher inmanganese nodules than in the adjacent sediment.

604

221. Bender, M.L., T.L. Ku and W.S. Broecker. 1966. Manganese
nodules: their evolution. Science VIS1, p. 325-328.
Arguments in favor of the very low growth rates foundfor manganese nodules on the basis of 230Th measurements.

222. Ku, T.L.

1966. Uranium series disequilibrium in deep sea
sediments. Ph.D. thesis. Columbia Uni~1~sit~~ New York, N.Y.
Detailed measurement of the 232 Th /U,
4U/238 U,
230 Th /234 U and 226 Ra /230 Th ratios in a series of deepsea cores.

223. Somayajulu, B.L.K. and E.D. Goldberg.

1966. Thorium anduranium isotopes in sea water and sediments. Earth & Planet.
Sci. Lett. VI, p. 102-106.
Discussion of thorium and uranium isotopes in seawaterand sediments.

224. KU, T.L. and W.S. Broecker, 1967.
tactinium in a manganese nodule.

V2, p. 317 - 320.

Uranium, thorium, proEarth & Planet. Sci. Lett.

Growth rates for a manganese nodule based on the distributions of 230 Th, 231 pa and 234 U in thinly sliced
samples.

225. Barnes, S.S. and J.R. Dymond.

1967. Rates of accumulationof ferromanganese nodules. Nature, V213, p. 1218-1219.
Measurements of the 4o Ar /4o K for the volcanic coresfound in manganese nodules and the growth rates obtained
from the 40 Ar age.

226. Ku, T.L., W.S. Broecker and N. Opdyke.

1968. Comparison of
sedimentation rates measured by paleo-magnetic and the ionium
methods of age determination. Earth & Planet. Sci. Lett. V4,
p. 1- 6.
Comparison of sedimentation rates over the last 300,000
years based on 230 Th measurements with those based on thedepth of the last magnetic reversal (i.e., that at about
700,000 years ago).

227. Ku, T.L. and W.S. Broecker.

1969. Radiochemical studies on
manganese nodules of deep-sea origin. Deep Sea Res. V16, p.
625-637.

Distributions of U and Th in nodules from different
depths in the sea.

228.

229.

Bender, M.L., T.L. Ku and W.S. Broecker. 1970. Accumulation
rates of manganese in pelagic sediments and nodules. Earth& Planet. Sci. Lett. V8, p. 143-148.
Comparison of the accumulation rates for Mn in nodules
and in pelagic sediments showing that the rate for sediments is higher than that for nodules.
Broecker, W.S., and J. Van Donk. 1976. Insolation
changes, ice volumes, and the 18 0 record in deep sea

605

cores. Reviews of Geophysics and Space Physics, V8, p. 169-196.
Revision of the Emiliani chronology for the last
several climatic cycles based on detailed 231 pa and
230 Th measurements on deep sea sediment core V12-122
from the C:aribbean Sea ..
230. Bender, M.L., W.S. Broecker, V. Gornitz, V. Middel, R. Kay,
S.S. Sun and P.E. Biscaye. 1971. Geochemistry of threecores from the East Pacific Rise. Earth & Planet. Sci.
Lett. V12, p. 425-433.
Manganese accumulation rates for ridge crest sediments
receiving the produ~~s of hydrothermal activity as
determined by the 2 Th method.
231. Cherdyntsen, V.V. 1971. Uranium-234. Israel Program for
Scientific Translation. Jerusalem, p. 1-2~4.
Discovery of the separation between 2 4 U and 238 U innature.
232. Somayajulu B.L.K., G.R. Heath, T.C. Moore, and D.S. Cronan.
1971. Rates of accumulation of manganese nodules andassociated sediments from the equatorial Pacific. Geochim.
Cosmochim. Acta, V35, ~. 621-624.
Measurements of 30 Th in manganese nodules.
233. Ku, T.L., J.L. Bischoff and A. Boersma. 1972. Age studies
of Mid-Atlantic Ridge sediments near 42°N and 20 o N. Deep Sea
Res. V19, p. 233-247.
Uranium series isotope measurements in deep sea coresfrom the flank of the mid-Atlantic ridge.
234. ThomsDn, J. and A. Walton. 1972. Natural radioactive decay
series elements in the oceans and sediments. Proc. R. Soc.
Edinb. Sere B. V15, p. 167-182.
Measurements of 231 pa and 230 Th in an equatorial
Atlantic core.
235. W.S. Broecker and S. Broecker. 1974. Carbonate dissolution on
the western flank of the east Pacific Rise. In: "Studies in
PaleQ-oceanography", edited by W.W. Hay, Society of Economic
Paleontologists and Mineralogists Special Publication No. 20,
pp.

4~3~±h measurements on cores from the East Pacific Rise.

236. Ku, T.L. 1976. The uranium-series methods of age determination. Ann. Rev. Earth & Planet. Sci. V4, p. 347-379.
A review of the uranium series dating methods asapplied to deep sea cores and manganese nodules.
237. Mangini, A. and C. Sonntag. 1977. 231pa dating of deep sea
cores via 227 Th counting. Earth & Planet. Sci. Lett. V37, p.
252-256.
An alternate means of measuring 231 pa in deep sea cores.

606

238. DeMaster, D.J. 1919. The marine budgets of silica and328i •
Ph.D. thesis, Yale University, New Haven, Conn.
Demonstration that cores taken beneath upwelling areas
have unusually high 231 pa /230 Th ratios.
239. Ku, T.L. 1979. Rates of manganese accretion. In: Marine
Manganese Deposits, ed. by G.P. Glasby, Elsevier, Amsterdam.
p. 249-268.
An excellent review paper on measurements of thegrowth rates of manganese nodules.
239a.Bacon, M.P. and J.N. Rosholt. 1982. Accumulation rates of
230 Th and 231 pa and some transition metals on the Bermuda
Rise. Geochim. Cosmochim. Acta V46, p. 651-666.
Demonstration that a dee~ sea sediment accumulating at avery high rate (36 cm/10 yrs) carries concentrations of
230Th and transition metals similar to sediments accumulating at normal rates (i.e., a few centimeters per
thousand years) and the implications of this finding toparticulate scavenging models.
239b.Veeh, H.H. 1982. Concordant 230 Th and 231 pa ages of marine
phosphorites. Earth & Planet. Sci. Lett. V51, p. 278-284.
Demonstration that reliable uranium series ages can beobtained on marine phosporites.

3. PAPERS DEALING WITH 210 Pb AND OTHER BIOTURBATIONTRACERS IN DEEP SEA SEDIMENTS (#240 to 6247)
240. Goldberg, E.D. and M. Koide. 1962. Geochronological studies
of deep sea sediments by the thorium and ionium method. Geochim;_ Cosmochim. Acta V26, p. 417-450.
First mention of the role of bioturbation in altering
the distribution of radioisotopes in marine sediments.
241. Berger, W.H. and Heath, G.R. 1968. Vertical mixing in pelagic sediments. Jour. Mar. Res. V26, p. 134-143.
A simple model for bioturbation using the distributionof volcanic ash as an example.
242. Guinqsso, N.L. and D.R. Schink. 1975. Quantitative estimates' of biological mixing rates in abyssal sediments.
Jour .. of Geophys. Res. V80, p. 3032-3043.
A modification of the Berger-Heath model to incorporatea finite rate of mixing within the bioturbated layer.
243. Nozaki, Y., J.K. Cochran, K.K. Turekian and G. Keller. 1977.
Radiocarbon and 210 Pb distribution in submersible-taken deepsea cores from Project FAMOUS. Earth & Planet. Sci. Lett.
V34, p. 167-173.
First demonstration that the excess 210 Pb in deep sea
cores could be used as a measure of the rate ofbioturbation.

607

244. SChink, D.R., and N.L. GUinasso, Jr.

1978. Redistributionof dissolved and absorbed materials in abyssal marine sediments undergoing biological stirring. Amer. Jour. Sci. V278,
p. 687-702.
Application of their bioturbation model to a series ofsubstances.

245. Peng, T.-H. W.S. Broecker and W.H. Berger.

1979. Rates ofbenthic mixing in deep sea sediments as determined byradioactive tracers. Quaternary Res. VII, P. 141-149.
Profile of excess 210 Pb in a deep sea core from thewestern equatorial Pacific Ocean.

246. Kadko, D. 1980. 23 °Th, 22 6Ra and 22 2 Rn in abyssal sediment s.
Earth & Planet. Sci. Lett. V49, p. 360-380.
Detailed analyses of all the key uranium series
isotopes (including 210 Pb ) in the upper 25 em of two
deep sea cores showing very low deposition rates.

247. Cochran, J.K. and S. Krishnaswami.

1980. Radium, thorium,
uranium and 210 Pb in deep-sea sediments and sediment porewaters from the North Equatorial Pacific. Amer. Jour. Sci.
V280, p. 849-889.
Measurements of 210 Pb in the pore waters of deep sea
sediments.

247a.Olsen, C.R., H.J. Simpson, T.-H. Peng, R.F. Bopp and R.M.
Trier. 1981. Sediment mixing and accumulation rate effects
on radionuclide depth profiles in Hudson estuary sediments.
J. Geophys. Res. V86 p. 11020-11028.
Use of 1 3 4 Cs , 1 ~ 7 Cs and 239, 24 °pu as tracers of sediment mixing and accumulation in estuaries.
-4. PAPERS DEALING WITH THE DISTRIBUTIONS OF 26 A1 , 32 S1 ,AND lOBe IN SEDIMENTS AND MANGANESE (#248 to #264)
24e. Goel, P.S., D.P. Kharkar, D. Lal, N. Narsappaya, B. Peters
and V. Yatirajam. 1957. The beryllium-10 concentration in
deep sea sediments. Deep Sea Res. V4, p. 202-210.
The first measurement of lOBe in deep sea sediments.
249. Merrill, J.R., E.F.X. Lyden, M. Honda, and J.R. Arnold.
1960;. The sedimentary geochemistry of the beryllium isotopes.
Geochim. et Cosmoshim. Acta, V18, p. 108-129.
A discussion of dating marine sediments with lOBe and
26 AI . The sedimentation rates determined from lOBe measurement are all consistent with those obtained from 14 C
dating.
250. Amin~ B.S., D.P. Kharkar, and D. Lal. 1966. Cosmogenic lOBe
and 6 Al in marine sediments. Deep-Sea Res. V13, p.
805-824.
A discussion of dating marine sediments with 1 DBe and26Al. The sedimentation rates determined from lOBe

608

measurement are all consistent with those obtained from
14C- dating.
251. Somayajulu, B.L.K. 1967. lOBe in a manganese nodule. Science
V156 p. 1219-1220.
i DEe dating of a manganese nodule as a confirmation of
their low growth rates.
252. Somayajulu, B.L.K. 1969. Study of marine processes using
naturally occurring radioactive nuclides. Ph.D. Thesis,
Univ. of Bombay, India.
Measurements of 32 Si in sea water.
253. Krishnaswami, S., B.L.K. Sornayajulu and W.S. Moore. 1972.
Dating of manganese nodules using beryllium-10. In: "Ferromanganese deposits on the ocean floor". ed., D.R. Horn,
Lamont-Doherty Geological Observatory, Palisades, N.Y. p.
117-122.
C~nfirmation of the mean growth rate for manganese
nodules by lOBe measurements.
254. Bhat, S.G., S. Krishnaswami, D. Lal, M. Rama and B.L.K.
Somayajulu. 1973. Radiometric and trace element studies offerromanganese nodules. In: Proc. Sump. Hydrogeochemistry
and Biogeochemistry 1~ The Clark Co. Wash., D.C., p. 443-462.
Measurement of 1 Be and 230 Th in ferromanganese nodules
from the Pacific Ocean.
255. Reyss, J.L. and Y. Yokoyama. 1976. Aluminum-26 in deep sea
sediments. Science V1~3, p. 1119-1121.
Measurements of 2 Al in an Al 2 03 sample prepared from a-Pacific sediment.
256. Guichard, F. J-L Reyss and Y. YOko~ama.
19~~.
Growth rateof manganese nodule measured with DBe and
AI. Nature
V272, p. 155-156.
Confirmation of radioactive decay hypothesis (as opposedto the inward diffusion hypothesis) for radioisotopes inmanganese nodule using lOBe and 26 AI.
257 • Raisbeck, G.M., F. Yiou, M. Fruneau, J.M. Soiseaux. 1978.
Beryilium-10 mass spectrometry with a cyclotron. Science
V202,·p. 215-217.
Description of a technique for the measurement of lOBe
using cyclotron mass spectrometry.
258. Boudreau, B.P. and M.R. Scott. 1978. A model for the diffusion-controlled growth of deep sea manganese nodules. Am. J.
Science, V278, p. 903-929.
Calculations based on hydrodynamic theory showing themagnitude of the resistance to chemical transport posedby the benthic boundary layer. Also it is shown that
this resistance could well be the limiting one formanganese nodule growth.

609

259. Raisbeck, G.M., ~o Yiou, M. Fruneau, J.M. Loiseaux, M.
Lieuvin. 1979.
Be concentration and residence time in theocean surface layer. Earth Planet. Sci. Lett. V43 p. 237240.
'
Measurements of lOBe concentration in Pacific surface
water by using the accelerator technique.
260. Ku, T.L., A. Omura and P.S. Chen. 1979. lOBe and U-series
isotopes in manganese nodules from ther central North Pacific. Marine Geology and Oceanography of the Pacific manganese
nodule province ed. by J.L. Bischoff and D.Z. Piper, Plenum
PUblishing Corp, p. 791-814.
lOBe and uranium series measurements on the same nodule
permitting a direct comparison of growth rates obtainedby these two methods. Diffusive redistribution of theseisotopes is suggested.
261. Turekian, K.K., J.K. Cochran, S. Krishnaswami, W.A. San{~rd,
P.D. Parker and K.A. Bauer. 1979. The measurement ofBe
in manganese nodules using a tandem Van de Graff accelerator.
Geophys. Res. Lett. V6, p. 417-420.
The first lOBe dating of nodules carried out using thehigh energy mass spectrometry method.
263. Sharma, P. and B.L.K. Somayajulu. 1980. Growth rate and composition of two ferromanganese nodules from the central North
Pacific. In: Sur la genese des nodules de manganese(Colloque international du CNRS No. 28q) p. 281-288.
Measurements of lOBe " U 230Th , 2j~Th , 226 Ra , 2l0 Pb " Mn
Fe in two ferromanganese nodules from the Central North
Pacific.
264. Krishnaswami, S., J.K. Cochran, K.K. Turekian and S.S. Sarin.
1980. -Time scales of deep-sea ferromanganese nodule growth
based on lOBe and a track distributions and their relation touranium decay series measurements in Sur la genese desnodules de manganese (C~lloque International du CNRS No.
289) p. 251-260.
Measurements of U/Th series nuclides, alpha track
density and lOBe to determine growth rates and decipher
growth history in ferromanganese nodules from thenortheast equatorial Pacific.
264a.Krishnaswami, S., A. Mangini, J.Ho Thomas, P. Sharma, J.K.
C.ochran, K.K. Turekian and P.D. Parker. 1982. lOBe and Th
isotopes in manganese nodules and adjacent sediments: nodule
growth histories and nuclide behavior. Earth & Planet. Sci.
Lett. V59, p. 217-234.
Demonstration that both Th and Be accumulate
preferentially in sediment over Mn nodules.

610

II. PAPERS DEALING WITH STABLE ISOTOPE TRACERS (#265 to #345)
A. PAPERS DEALING WITH STABLE ISOTOPE TRACERS OF WATER MOTION(6265 to 284)
1. PAPERS DEALING WITH THE DISTRIBUTION OF THE 3He/ 4 HeRATIO IN THE SEA (#265 to #281)
265. Suess, H.E. and H. Wanke.
1965. On the possibility of ahelium flux through the ocean floor. In:
Progress in
Oceanography, V3, Pergamon Press, Oxford, p. 347-353.
Assuming that the 4He escape from the earth's interiormatches its production by the decay in the earth's
interior, a prediction of the escape rate of this gas
into the deep sea is made.
Assuming that as suggested
by 14 C data the residence time of water in the deep
Pacific is 700 years, an excess of 20% in the 4 He
content of deep Pacific Ocean water is predicted.
Measurements are reported showing that the actual
excess is not more than 6%.
266. Bieri, R.H., M. Koide and E.D. Goldberg. 1964.
Noble gases
in sea water. Science, V146, p. 1035-1037.
An attempt to measure the excess 4He in sea water.
The
values reported have subsequently been shown to be inerror.
267. Bieri, R.H., M. Koide and E.D. Goldberg. 1966. The noble gas
contents of Pacific seawaters.
J. of Geophys. Res. V71, p.
5243-5265.
Supersaturation of noble gases including He, Ne, Ar, andKy in seawaters is discovered.
The distributions ofthese gases with depth and their relationship to mixing
and advection are discussed.
267a.Bieri, R.H., M. Koide and E.D. Goldberg. 1968. Noble Gas
C~ntents of Marine Waters.
Earth & Planet. Sci. Lett. V4,
p. 329-340.
Measurements of He, Ne, Ar, and Kr in the Atlantic andPacific Oceans showing deviations from saturation.
268. Mamyrin, B.A., I.N. Tolstik1ien, G.S. Anufriev, and I.L.
Kamenski.
1969.
Anomalous isotopic composition of helium involcanic gases.
Dokl. Akad. Nank.
SSSR, V184, p. 1197.
The discovery of primordial 3 He in volcanic gases.
269. Clarke, W.B., M.A. Beg and H. Craig.
1969. Excess 3 He in
the Atlantic.
Earth & Planet. Sci. Lett., V6, p. 213-220.
The discovery of excess 3 He in sea water and thedemonstration that this 3 He must be primordial.
1970. Excess helium-3
270. Clarke, W.B., M.A. Beg and H. Craig.
at the North Pacific GEOSECS station. Jour. of Geophys. Res.

611

Res. V75, p. 7676-7678.
Profile of 3He /4 He ratio for a station in the eastern
North Pacific. The maximum in this ratio at a depth ofabout 3 kilometers depth bears witness to the release of
3He from the ridge crests.
271. Jenkins, W.J.,M.A. Be~, W.B. Clarke, P.J. Wangersky, and H.
Craig. 1972. Excess He in the Atlantic Ocean. Earth &Planet. Sc. Lett. V16,p.122-126.
First results on excess 3He in waters from theAtlantic Ocean.
272. Craig, H., W.B. Clarke and M.A. Beg. 1975. Excess 3 He in
deep water on the East Pacific Rise. Earth and Planet. Sci.
Lett. V26, p. 125-132.
Measurements of excess 3 He in the Pacific Ocean close
to the ridge crest showing that the anomalies are, as
expected, higher close to the source of mantle exhalations.
273. Lupton, J.E. 1976. The 3He distribution in deep water over
the Mid-Atlantic Ridge. Earth and Planet. Sci. Lett. V32, p.
371-374.
Measurements showing the absence of a 3He anomaly over
the Mid-Atlantic Ridge. Evidence for an excess in 4He
in the deep eastern Atlantic is also given.
274. Jenkins, W.J. and W.B. Clarke. 1976. The distribution of
3He in the western Atlantic Ocean. Deep Sea Res. V23, p.
481-494.
The distribution of 3 He excesses in the western Atlan_tic. In the northern part of this basin these anomalies
are largely from the decay of tritium. In the southernpart deep water anomalies are generated by the entry
from the Antarctic of water previously contaminated withprimordial helium in the Pacific Ocean.
275. Jenkins, W.J. 1977. Tritium-helium dating in the Sargasso
Sea: A measurement of oxygen utilization rates. Science
V196, p.291-292.
Correlation of the excesses of 3He produced by tritium"decay with oxygen deficiencies in the thermocline of thetemperate North Atlantic.
276. Jenkins, W.J., J.M. Edmond, and J.B. Corliss. 1978. Excess
3He and 4He in Galapagos submarine hydrothermal water. Nature
V272, p. 156-158.
The relationship between the release of 3He and of heatfrom the mid-ocean ridges. Also these measurements
yield the ratio of excess 3 He to excess 4He entering thesea from ridge crests.
277. Jenkins, W.J., W.V. Collentro and R.D. Boudreau. 1979. Woods

612

Hole Oceanographic Institution. Helium Isotope Laboratory
Data Release #1.
Listing of tritium and 3 He data from the Sargasso Sea.
278. Broecker, W.S. 1980. The distribution of 3 He anomalies inthe deep Atlantic. Earth and Planet. Sci. Lett. V49, p.
513-519.
A discussion of pUblished data in which an attempt ismade to separate the contributions of excess 3 He added
directl~ to the Atlantic from the contribution ofexcess He carried into the Atlantic by waters fromthe Antarctic (previously contaminated by ridge crest
emissions in the Pacific). Evidence is also given fora narrow plume of excess 3He bearing water hugging thecontinental margin of North America. The origin of
this plume is not known.
279. Jenkins, W.J. 1980. Tritium and 3He in the Sargasso Sea. Sears
Foundation: J. of Marine Research, V38 p. 533-569.
Vertical profiles of tritium and ~He are shown to be incompatible with one-dimensional mixing models. Alongisopycnal ventilation of the temperate thermocline must
be invoked if these distributions are to be explained.
280. Lupton, J.E. and H. Craig. 1981. A major helium-3 source at
15°3 on the East Pacific Rise. Science V214, p. 13-18.
3 He excesses of 50% are reported over the East Pacific"Rise. A plume of excess helium at 2.5 km depth isshown to extend 2000 krn to the west of the rise crest.
281. Craig, H. and J.E. Lupton. 1981. Helium-3 and Mantle Volatile& in the Ocean and the Oceanic Crust. The Sea V7: The
Oceanic Lithosphere. Edited by C. Emiliani, Wiley and Son,
New York, p. 391-428.
An excellent review of research dealing with the releaseof rare gases by the oceanic crust and the movement ofthese gases through the ocean.
2. PAPERS DEALING WITH THE DISTRffiUTIONS OF D/H AND OF
18 0 / 16 0 RATIOS IN THE OCEAN (1282 to '284)

282. Epstein, Sand T. Mayeda. 1953. Variation of 18 0 content ofwaters from natural sources. Geochim. Cosmochim. Acta V4,
p. 213-224.
First paper showing the variations in the 18 0/16 0
ratio within the oceans.
283. Craig, H. and L.I. Gordon. 1965. Isotopic Oceanography:
Deuterium and oxygen-18 variations in the ocean and themarine atmosphere. Proc. Symp. Mar. Geochem., Univ. of Rhode
Island, Occ. Publ. 3, p. 277-374.
A discussion of the origin of bottom waters formed in
the Antarctic region based on the 18 0 /16 0 ratios in613

various oceanic water types.
284. Weiss, R.F., H.G. Ostlund and H. Craig.
1979. Geochemical
studies in the Weddell Sea. Deep Sea Res. V26, p. 1093-1120.
Use of 18 0 /16 0 and D/H ratios to aid in defining thesource of Weddell Sea Bottom Water.

B. PAPERS DEALING WITH STABLE ISOTOPE TRACERS OF BIOLOGICALPROCESSES IN THE SEA (#285 to #317)
1. PAPERS DEALING WITH THE 13 Cj12C RATIO IN DISSOLVED
OCEANIC CARBON AND IN MARINE ORGANISMS (#285 to

1305)
285. Craig, H. 1953. The geochemistry of the stable carbon isotopes.
Geochim. Cnsmochim. Acta. V3, p. 53-q2.
Paper dealing the distribution of 13 C/12 C ratios innatural materials including marine macrophytes andalgae.
286. Keeling, C.D.
1961. The concentrations and isotopic abundances of carbon dioxide in rural and marine air.
Geochim.
Cosmochim. Acta, V24, p. 277-298.
C02 partial pressures and 13 C/12 C ratios for atmospheric
CO 2.
287. Sackett, W. M., W. R. Eckelmann, M. C. Bender and A. W. H.
Be.
1965. Temperature dependence of carbon isotope composition in marine plankton and sediments. Science V148, p. 235237.
13 C;12C ratios in marine plankton from various parts of
the ocean.
288. Deuser, W. G. and E. T. Degens.
1~67.
Carbon isotope fractionation in the system C02 (gas) - C02 (aqueous). Nature
V215, p.1033-1035.
Measurements of the carbon isotope fractionation factor
between atmospheric C02 and surface ocean LC02.
289. Wendt, I.
1968. Fractionation of carbon isotopes and its
temperature dependence in the system C02 gas-C02 in solution
and HC03--C:02 in solution. Earth & Planet. Sci. Lett. V4, p.
64.
Measurements of the carbon isotope fractionation factor
between gas phase CO 2 and dissolved CO 2 and HC0 3 -.
290. Williams, P. M. and L. I. Gordon.
1970. Carbon-13:carbon-12 ratios in dissolved and particulate organic matterin the sea.
Deep Sea Res. V17, p. 19-27.
13 C/12 C ratios in particulate carbon filtered fromsea water.
291. Vogel, J.C., P.M. Grootes and W.G. Mook. 1970. Isotopic fractionation between gaseous and dissolved carbon dioxide. Z.
Physik V230, p. 225-238.

614

Measurements of the carbon isotope fractionation factor
between gas phase and aqueous phase CO 2 .
292. Kroopnick, P., W. G. Denser and H. Craig. 1970. Carbon-13
measurements on dissolved inorganic carbon at the North
Pacific (1969) GEOSECS station. J. Geophys. Res. V75 p.
7668-1 671 •',3C/12C ratios in oceanic L:CO 2.
293. Craig, H. 1970. Abyssal carbon-13 in the South Pacific.
J. Geo~hys. Res., V75, p. 691-695.
1 C/ 12C ratios in oceanic ECO 2 .
294. ~~oopnick,.p., R. F. Weiss aq~ H. Craig. 1972. Total C02,
C., and dlssol ved oxygen a at GEOSECS II in the North
Atlantic. Earth Planet. Sci. Lett., V16, p. 103-110.
13 C/12 C ratios in oceanic L:C02.
295. Kroopnick, P. 1974. Correlations between 13 C and EC02 insurface waters and atmospheric C02. Earth Planet. Sci.
Lett. J V22~ p. 397-403.
13 C/ L~C ratios in oceanic ECO 2.
296. Kroopnick, P. 1974. Modeling the dissolved 02 - C02 _ 13 C
system in the eastern equatorial Pacific. Deep Sea Res. V21,
p. 21l-22~.
13 C/ 2C ratios in oceanic C02.
297. Mook, W.G., J.C. Bomrnerson and W.H. Stoverman. 1974. Carbon
isotope fractionation between dissolved bicarbonate and gaseous carbon dioxide. Earth Planet. Sci. Lett. V22, p. 69-176.
Measurements of the carbon isotope fractionation factor-between atmospheric C02 and dissolved gaseous carbon andbicarbonate ion.
13C var<5298. Kroopnick, P., s. V. Margolis and C.S .. Wong. 1977.
variations in marine carbonate as indicators of the C02
balance between the atmosphere and oceans. In: The Fate ofFossil Fuel C02 in the Oceans, N.R. Anderson and A. Malahoff,eds., Plenum Press, New York, N.Y.,p.295-231.
13 C/12 C ratios in oceanic EC02.
299. Siegertthaler, U., M. Heimann and H. Oeschger. 1~78. Model
responses to the atmospheric C02 level and 13C/ l C ratio tobiogenic C02 input. In: Williams, J. Ced.), Carbon dioxide,climate and society, Pergamon Press, New York, p. 79-87.
Calculations using the Oeschger model showing the impactof releases of soil and living biosphere C02 during thepast century on the 013 C record for atmospheric C02.
300. Keeling\ C.D., M.G. Mook and P.P. Tans. 1979. Recent trend
in the 3C/12 C ratio of atmospheric carbon dioxide. Nature,
V277, p. 121-123.
Update on the 13 C/12 C ratio in atmospheric C02. Taken

615

togethe~ with the ea~lier Keeling results, an estimate
can be made of the rate of decline of this ~atio due tothe burning of fossil fuels and of changes in the living
biomass and soil carbon reservoi~.

1980. The dist~ibution of 13 C in the Atlantic
Ocean. Earth Planet. Sci. Lett., V49, p. 469-484.
13C/12C ratios for oceanic EC0 2 as obtained for the
Atlantic GEOSECS Program.

301. Kroopnick, P.

302. Keeling, C.D. R.B. Bacastow and P.P. Tans. 1980. Predicted
shift in the l 3C/12C ratio of atmospheric carbon dioxide.
Geophys. Res. Lett., V7, p. 505-508.
If a carbon isotope kinetic separation factor of 0.986
for the C02 entering the sea is assumed, these authors
show that the time history of the o13 C for atmosphericC02 does not provide a very sensitive way to estimate
the contribution of forests and soils to the anthropogenic build up of the atmospheric C02 concentration.
However, if the 0.998 value of Seigenthaler and Munnich
is used then the situation is much improved.
Ostlund, H.G.
1980. GEOSECS Atlantic and Pacific 13 C meassurements on radiocarbon samples. Tritium Laboratory Data
Release #80-09, Rosenstiel School of Marine and Atmospheric
Scienc s, University of Miami.
11C/ 12 C ratios obtained on the CO 2 collected forradiocarbon analyses of the Atlantic and Pacific
GEOSECS program.

304. Siegenthaler, U. and K.O. Munnich, 1981.

13C fractionation
during C02 transfer from air to sea.
In: SCOPE report #16.
Carbon Cycle Modeling, B. Bolin, ed., John Wiley & Sons,

p.249-257.
Estimate of the kinetic fractionation during uptake ofC02 by the ocean.

2. PAPERS DEALING WITH THE 18 0/ 16 0 RATIOS IN ATMOSPHERIC
02 AND IN THE 02 DISSOLVED IN THE SEA (#306 to #317)
306. Dole, M.

1935. The relative atomic weight of oxygen in waterand a1r.
J. Am. Chern. Soc., V57, p. 2731.
The first in a lon~ series of papers dealing with thedifference in the 8 0 /16 0 ratio in atmospheric O2 and insea H 2 0. The measurements were made by density determination in the era before precise mass spectrometry.

307. Dole, M. and G. Jenks.

1944. Isotopic composition ofphotosynthetic oxygen. Science, v. 100, p. 409.
Another in the Dole series.

308. Dole, M., R.C. Hawkins and H.A. Barker. 1947- Bacterial fractionation of oxygen isotopes. J. Amer. Chern. Soc. V69, p.

616

226-22e.
Another in the Dole series.
309. Rakestraw, N.W., D.P. Rudd and M. Dole. 1951. Isotopic
composition of oxygen in air dissolved in Pacific Ocean water
as a function of depth. J. Am. Chern. Soc., V73, p. 2976.
Another in the Dole series.
310. Dole, M., B.A. Lane, D.P. Rudd and D.A. Zankelies. 1954.
Isotopic composition of atmospheric oxygen and nitrogen.
Geochim. Cosmochim. Acta, V6, p. 65-7e.
Another in the Dole series.
311. Lane, G.A. and M. Dole. 1956. Fractionation of oxygenisotopes during respiration. Science V123, p. 574-576.
Another in the Dole series.
312. Kroopnick, P. and H. Craig. 1971. Isotopic composition ofmolecular oxygen in the atmosphere and the sea. Trans. Am.
Geophys. Union V52, p. 255.
Measurements of the isotopic composition of O2 fromsea water.
313. Kroopnick, P. and H. Craig. 1972. Atmospheric oxygen:isotopic composition and solUbility fractionation. Science,
V175, ~. ~4-56.
1 0/ 6 0 fractionation between gaseous and dissolved 02.
314. Kroopnick, P., R.F. Weiss and H. Craig. 1972. Total C02,
13 C, and dissolved oxygen-18 at GEOSECS II in the North
Atlantic. Eatth Planet. Sci. Lett. V16, p. 103-110.
An 180/1
profile for the 02 dissolved in the sea at a-station in the eastern North Pacific.

°

315. Kroopnick, P. 1975. Respiration, photosynthesis, andoxygen isotope fractionation in oceanic surface water.
Limnot. 01~anogr. V20, p. 988-992.
80/
measurements on 02 dissolved in the sea.

°

316. Kroopnick, P. and H. Craig. 1976. Oxygen isotope fractionation in dissolved oxygen in the deep sea. Earth Planet. Sci.
Lett .. V32 p. 375-378.
180/~60 measurements on 02 dissolved in seawater.
317. Kroopnick, P. 1980. Isotopic fractionations during oxygen
consumption and carbonate dissolution within the North
Atlantic deep water. Earth Planet. Sci. Lett. V49, p. 485498.
18 0 /16 0 measurements on dissolved 02 samples collectedduring the Atlantic GEOSECS expedition.

c. STABLE ISOTOPE TRACERS IN SEDIMENTS AND MANGANESE NODULES
1318 to #345)

617

1. PAPERS DEALING WITH THE 18 0 /16 0 RATIOS IN CaCO 3 FROMDEEP SEDIMENT ('318 to 6326)

318. Emiliani, C. 1955. Pleistocene temperatures. J. Geol.
V63, p. 538-578.
The first paper dealing with the 18 0/16 0 method as applied to deep sea cores. The depiction of climatic
cycles from the down core records is a classic piece of
science which had a great impact on the field of paleoclimatology.
Broecker, W.S. and J. Van Donk. 1970. Insolation changes
ice volumes and the 18 0 record in deep-sea cores. Rev.
'
Geophys. Space Phys. V8, po 169-196.
A demonstration of the relationship between the 18 0 /16 0
record in deep-sea cores and the sea level record as
represented by the flight of uplifted coral reefs on theisland of Barbados.
320. Shackleton, N.J. and N.D. Opdyke. 1973. Oxygen isotope andpaleomagnetic stratigraphy of Equatorial Pacific core V28238: oxygen isotope temperatures and ice volumes on a
10 5 -year and 10 6 -year scale. Quat. Res., v. 3, ~. i~-55.
The first long and detailed record of the 1 0/ 0 ratio
in both planktonic and benthic forams.
321. Shackleton, N.J. 1974. Attainment of isotopic equilibrium
between ocean water and benthonic foraminifera genus
UVigerina:
isotopic changes in the ocean during the lastglacial. Colloques Internationaux du C~ntre National de laRecherche Scientifique, V219, p. 203-210.
The record of 8 0 /16 0 , as preserved in benthic forams.
~As the temperature of the deep sea cannot have changed
much between glacial and interglacial time, this record
provides an estimate of the 18 0 /16 0 ratio changes in thesea as a whole caused by the growth and retreat of continental ice sheets.
322. Shackleton, N.J. and N.D. Opdyke. 1976. Oxygen isotope andpaleomagnetic stratigraphy of Equatorial Pacific core V28239, late Pliocene to latest Pleistocene. In: Investigationof Late Quaternary Paleoceanography and Paleoclimatology, ed.
R. M.'Cline and J. D. Hays, Geol. Soc. Amer. Mem. V145, p.
449-464.
An even longer detailed record of the 18 0/ 16 0 variations
in planktonic and benthic forams.
Hays, J.D., J. Imbrie, N.J. Shackleton. 1976. Variations in
the earth's orbit: pacemaker of the ice ages. Science, V194,
p. 1121-1132.
Demonstration that the 18 0/ 16 0 records in deep sea cores
have periodicities of about 20,000, 40,000 and 100,000
years. As two of these (the 20,000 and 40,000) match
those in the earth's insolation record, this work pro-

618

324.

vides strong support for the influence of changes in theearth's orbital characteristics on climate.
Shackleton, N.J. 1977. The oxygen isotope record of the late
Pleistocene: Phil. Trans. Roy. Soc. Land. B, V280, p.

169-182.

A review of the 180 /16 0 record from deep-sea cores.

325. Duplessy, J.-C. 1978. Isotope studies. In: Climatic Change,
ed. John Gribbin, Cambridge Press, p. 46-47.
A review article on the geochemistry of 18 0 in the ocean
and atmosphere ..

326. Broecker, W.S. 1978. The cause of glacial to interglacial
climatic change.
In:"Evolution of Planetary Atmosphere andClimatology of the Earth", Centre National d'Etudes Spatiales(France), p. 165-177.
An apparent inconsistency between the amplitude of ice
volume changes as recorded by 18 0 /16 0 in benthic forams
and by raised coral reefs.

2. PAPERS DEALING WITH THE 13C/12C RECORD FOUND IN
FORAMINIFERA FROM DEEP SEA SEDIMENTS (#321 to #331)

327. Van Donk, J. 1970.

The Oxygen Isotope Record in Deep Sea
Sediments.
Ph.D. Thesis. Columbia University, New York City.
The first published detailed down core records of
13 C/12 C ratio in planktonic foraminifera.

328. Shackleton, N.J.

1977. Tropical rainforest history and theequatorial Pacific carbonate dissolution cycles.
In: TheFate of Fossil Fuel CD2 in the Oceans, eds. N.R. Anderson and
A. Malahoff, Plenum Press, New York p. 401-428.
:A demonstration that the 13C/l~C ratio in the dissolved
inorganic carbon from abyssal waters in the ocean underwent glacial to interglacial changes and the implicationof these changes to the earth's carbon budget. A classic paper!

329. Berger, W.H. and J.S. Killingsley.

1977. Glacial-Holocenetransition in deep-sea carbonates:
selective dissolution andthe stable isotope signal.
Science, v. 197, p. 563-566.
Demonstration that the 13C/12 C ratio in planktonic'organisms underwent almost no change from glacial tointerglacial time.

330. Belanger, P.E., W.B. Curry and R.K. Matthews.

1981. Core-top
evaluation of benthic foraminiferal isotopic ratios forpaleo-oceanographic interpretations. Palaeogeog. Paleoclimat. Palaeoecol., V33, p. 205-220.
The 13C and 18 0 variations in benthic foraminifera areused as sensors of bottom water conditions.

331. Broecker, W.S.

1981. Glacial to interglacial changes inocean and atmosphere chemistry. In: Climatic Variations and619

Variability: Facts and Theories, A. Berger, editor, D. Reidel
Publishing C~o., 109-120.
An explanation for the difference between the 13C/ 12 C
record for planktonic and benthic foraminifera between
glacial and interglacial time.
331a.Curry, W.B. and R.K. Matthews. 1981. Equilibirum 18 0 fractionation in small size fraction planktic foraminifera:
Evidence from recent Indian Ocean sediments, Marine Micropaleontology, V6, p. 327-33I.
Measurements of both 0 80 and o13 C in different size
fractions of a single species of foraminifera showingthat the size dependence of the 13 C/12 C ratio is not
simply an artifact of changing depth habitat withmaturity.

3. PAPERS DEALING WITH 143 Nd /144 Nd RATIOS IN SEA WATER,
MARINE SEDIMENTS AND MANGANESE NODULES (#332 to #340)
332. O'Nions, R.K., S.R. Carter, R.S. Cohen, N.M. Evensen and
P.J. Hamilton. 1978. Pb, Nd and Sr isotopes in oceanic ferromanganese deposits and ocean floor basalts. Nature, V273,
p. 435-438
Measurements of 143 Nd /144 Nd ratios in manganese nodules
and in metalliferous sediments.
333. McCullock, M.T. and G.J. Wasserburg. 1978. Sm-Nd and Rb-Sr
chronology of continental crust formation. Science, V200,
p. 1003-1011.
Measurements of 143 Nd /144 Nd ratios in red clay.
334. Cohen, R.S., N.M. Evensen, P.J. Hamilton and R.K. O'Nions.
1980;. U-Pb, Sm, Nd, and Rb-Sr systematics of mid-ocean ridge
basalt glasses. Nature, V283, p. 149-153.
Measurements of Pb, Nd, and Sr isotopes in basalt
glasses from mid-ocean ridges.
335. O'Nions, R.Ke, N.M. Evensen and P.J. Hamilton. 1979. Geochemical modeling of mantle differentiation and crustal growth.
J. Geophys. Res., V84, p. 6091-6101.
A discussion of two-reservoir model with time-dependenttransport coefficient between the reservoirs to model
the abundances of K, Ar, Rb, Sr, Sm, Nd, U, Th, and Pb
and the isotopic compositions of Ar, Sr, Mn, and Pb in
the earth's mantle and continental crust.
336. Piepgras, D.J., G.J. Wasserburg, E.J. Dasch. 1979. The isotopic composition of Nd in different ocean masses.
Earth Planet. Sci. Lett., V4~, p. 223-236.
Measurements of 143Nd/144Nd in manganese nodules, metalliferous sediments, red clays, hydrothermal crusts andsea water.
337. Piepgras, D.J. and G. J. Wasserburg. 1980. Neodymium isotopic

620

variations in seawater. Earth Planet. Scio Lett. V50, p. 128-

138.

Measurements of 143Nd/144Nd ratios in seawater samples
from the Atlantic and Pacific Oceans.

338. Allegre, C.J. and Othman, B. 1980. Nd-Sr isotopic relationship in granitoid rocks and continental crust development: A
chemical approach to orogenesis. Nature, V286, p. 335-342.
A discussionof a mixing model to explain the generationof granites from the mantle and the recycling of continental material based on Nd-Sr isotopic compositions.

339. Goldstein, S.L. and R.K. O'Nions.

1981. Nd- and Sr-isotopic
relationsships in pelagic clays and ferromanganese nodules.
Nature, V292, p. 324-12§.
Measurements of 4 Nd/ 144 Nd ratios in marine sediments
and manganese nodules.

340. Taylor, S.R. and S.M. McLennan. 1981.

The composition andevaluation of the continental crust: rare earth element evidence from sedimentary rocks. Phil Trans. Roy. Soc. Lond. A,
V301, p. 381-399.
A discussion of the composition and evolution of thecontinental crust using rare earth element patterns fromsedimentary rocks.

4. PAPERS DEALING WITH LEAD ISOTOPE RATIOS IN DEEP SEA
SEDIMENTS AND MANGANESE NODULES (#341 to #345)

341. Chow, T.J. and C.C. Patterson. 1959. Lead Isotopes in manganese nodules. Geochimica et Cosmochimica Acta, V17, p. 21-31.

A discussion of isotopic composition of lead inmanganese nodules of the Atlantic and Pacific Oceans.

342. Chow, T. and C.C. Patterson.

1962. The occurrence and significance of lead isotopes in pelagic sediments". Geochim.
Cosmochim. Acta, V26, p. 263-308.
Lead isotope measurements in deep sea sediments andmanganese nodules.

343. Reynolds, P.H. and E.J. Dasch.

1971. Lead isotopes in marine manganese nodules and the ore-lead growth curve. J. Geop hy s. Re s ., v. 76, p. 5124- 512 9.
Lead-isotope composition in marine manganese nodules.

344. O'Nions, R. K., S.R. Carter, R.S. Cohen, N.M. Evensen and
P.J. Hamilton. 1978. Pb, Nd and Sr isotopes in oceanic
ferromanganese deposits and ocean floor basalts.

V273, p. 435-438.

Nature,Measurements of Pb-, Nd- and Sr-isotope compositons ofoceanic ferromangenese deposits.

621

345. Dasch, E.J. 1981. Lead isotopic composition of metalliferous
sediments from the Nazca plate. Geol. Soc. Am. Memoir 154,
p. 199-210.
Lead isotope measurements in hydrothermal deposits
compared to those in manganese nodules and basalts.

III. PAPERS DEALING WITH CHEMICAL TRACERS (#346 to #456)
A. PAPERS DEALING WITH CHEMICAL TRACERS OF BIOLOGICAL CYCLES('346-#433).

1. PAPERS DEALING WITH THE DISTRIBUTION OF N03, P0 4 , AND
DISSOLVED OXYGEN IN THE SEA ('346 to #364)
346. Redfield, A.C. 1934. On the proportions of organic
derivatives in sea water and their relation to thecomposition of plankton. James Johnson Memorial Vol.,
Liverpool, p. 176-192.
The relationship between the composition of marine
plants and animals and the variations in the N03, P0 4 ,~C02 and O2 content of sea water.
347. Redfield, A.C. 1942. The processes determining the concentrations of oxygen, phosphate and other organic derivatives
within the depths of the Atlantic Ocean. Papers Phys.
Oceanogr. Meteorol., V9, p. 1-22.
Again the relationship between the chemical gains andlosses during respiration and the chemical changes indeep water.
348. Rakestraw, N.W. 1947. Oxygen consumption in sea water overlong periods. J. Mar. Res. V6, p. 259-263.
-niscussion of the rate and extent of oxygen comusmption
in sea water under a variety of conditions.
349. Richards, F.A. and A.C.. Redfield. 1954. A correlation
between the oxygen content of sea water and the organic
content of marine sediments. Deep Sea Res., VI, p. 279-281.
The lower the oxygen content of sea water the higher
the organic carbon content of underlying sediments.
350. Riley, CoA. 1954. Oxygen, phosphate and nitrate in the
Atlantic Ocean. Bull. Bingham Oceanogr. ColI. 13-14, p. 1126.
A model for the distribution of respiration products inthe deep Atlantic.
351. Richards, F.A. 1957. Oxygen in the ocean. Memoirs Geolog.
Soc. Amer., v. 67, p. 185-23e.
Distribution of dissolved oxygen in the oceans.
352. Redfield, A.C.. 1960. The distribution of phosphorous in the
deep oceans of the world. Proc.-Verbaux Assoc. Oceanogr.
Phys. Union Geol. Geophys. Intern., v. 7, p. 189-193.

622

Distribution of P0 4 in the deep waters of the world
ocean.
353. Wyrtki, K. 1962. The oxygen minima in relation to ocean
circulation. Deep Sea Res., V9, p. 11-23.
Application of a one-dimensional advection-diffusion
model to the vertical distribution of dissolved oxygenin the ocean.
354. Redfield, A.C~., B .H. Ketchum and F .A. Richards. 1963. The
influence of organisms on the composition of sea water. In:
The Sea, V2, edited by M.N. Hill, Interscience, New York,
p. 26-77.
Review article on the relationship between the changes
in the chemical composition of sea water and the composition of marine organisms.
355. Carpenter, J.H. 1965. The Chesapeake Bay Institute technique for the Winkler dissolved oxygen method. Limnology andOceanography VI0, p.141-143.
Description of the universally used method for themeasurement of dissolved oxygen in sea water.
356. Gordon, A.L. 1966. Potential temperature, oxygen and circulation of bottom water in the Southern Ocean. Deep Sea Res.
V13, p. 1125-1138.
The relationship between the distribution of dissolved
oxygen and the hydrography of the Antarctic Ocean.
357. Menzel, D.W. and H.H. Ryther. 1968. Organic carbon and theoxygen minimum in the South Atlantic Ocean. Deep Sea Res.
V15, .p. 327-337.
A discussion of the relationship between the rain rateof organic matter and the rate of ventilation in thegeneration of oxygen minimum layers.
358. Weiss, R.F. 1970. The solubility of nitrogen, oxygen andargon in water and seawater. Deep-Sea Res. V17, p. 721-735
The relationship between gas solubility and temperatureand the salinity.
359.

K. 1971. Oceanographic Atlas of the International
Indian Ocean Expedition. National Science Foundation, Wash.

Wyrtk~,

D.C:.

Sections of the hydrographic and nutrient constituentsin the Indian Ocean.
360. Spencer, D.W. 1972. GEOSECS II, the North Atlantic station: hydrographic features of oxygen and nutrients. Earth& Planet. Sci. Lett. V16, p. 91-102.
Relationship between the nutrient constitutent profiles
and the profiles of potential temperature and salinity
at a station in western North Atlantic.

623

361. Broecker, W.S. 1974.
"NO" a conservative water-mass trace~.
Earth & Planet. Sci. Lett. V23, p. 100-107.
Demonst~ation that by adding the concent~ation of dissolved oxygen to nine times the nitrate concent~ationa quasi-conservative property is obtained. As this property has a far different value for deep waters formedin the Antarctic Ocean than for those formed in thenorthern Atlantic it provides a useful adjunct to theconventional 8-S method of distinguishing water types.
361a.Alvarez-Borrego, S., D. Guthrie, C.H. Culberson and P.K.
Park. 1975. Test of Redfield's model for oxygen-nutrient
relationships using regression analysis. Limnol. & Oceanog.
V20, p. 795-805.
Derivation of Redfield ratios from regression analysesof concent~ation data for dissolved constituents in thesea.
361b.Froelich, P.N., M.L. Bender, N.A. Luedtke, G.R. Heath and
T. DeVries.
1982. The marine phosphorus cycle, Am. J. Sci.,
V282, p. 474-511.
Discussion of sources and sinks of phosphorus in theocean.
The input and output flux and the mean residence
time for phosphorus in the ocean are given.
362. Bainbridge, A.E.
1975. GEOSECS Atlantic Ocean: final hydrographic data report.
GEOSECS Operations Group/National
Science Foundation, Wash. D.C.
Listing of the potential temperature, salinity, density,and nutrient constituent concent~ations for samples
obtained during the Atlantic Ocean GEOSECB expedition.
363.

Bainb~idge,

A.E.
1976. GEOSECB Pacific Ocean: final hydrographic data report. GEOSECS Ope~ations Group/National
Science Foundation, Wash. D.C.
Listing of the potential temperature, salinity, density,and nutrient constituent concentrations for samples
obtained during the Pacific Ocean GEOSECS expedition.

364. Bainbridge, A.E.
1978. GEOSECS Indian Ocean: final hydrographic data report.
GEOSECS Operations Group/National
Science Foundation, Wash. D.C.
Listing of the potential temperature, salinity, density,and nutrient constituent concentrations for samples
obtained during the Indian Ocean GEOSECB expedition.
364a.Williams, R.T. 1981. Hawaii-Tahiti Shuttle Experiment; Hydrographic Data Reports, VI, 2 and 3. Scripps Institution of
Oceanography (reference nos. 81-5, 6 and 7) Univ. of California, San Diego.
Nutrient constituent and oxygen for bimonthly traverses
from 20 0 S to 20 0 N in the central Pacific from Feb. 1979
to Feb. 1980.

624

2. PAPERS DEALING WITH THE DISTRIBUTIONS OF H4 Si0 4 , Ge
AND Ba IN THE SEA (#365 to 1378)
365. Wolgemuth, K. and W.S. Broecker. 1970. Barium in sea water.
Earth and Planet. Sci. Lett. V8, p. 372-378.
First demonstration of the close parallel between the
distributions of barium and 226 Ra in the sea.
366. Wolgemuth, K. 1970. Barium analyses from the first GEOSECB
test cruise. J. Geophys. Res. V75, p. 7686-7687.
A profile of barium in the eastern North Pacific.
367. Bernat, M., T. Church, C.J. Allegre. 1972. Barium andStrontium concentrations in Pacific and Mediterranean sea
water profiles by direct isotope dilution mass spectrometry,
Earth & Planet. Sci. Lett. V16, p.75-80.
A novel procedure for rapid barium analysis.
368. Bender, M., T. Snead, L.-H. Chan, M.P. Bacon and J. M.
Edmond. 1972. Barium intercalibration at GEOSECS I and III.
Earth & Planet. Sci. Lett. V16, p. 81-83.
Intercomparisons of barium analyses made in different
laboratories.
369. Bacon, M.P. and J.M. Edmond. 1972. Barium at GEOSECS III in
the southwest Pacific, Earth and Planet. Sci. Lett. V16, p.
66-74.
A profile of barium in the southwestern Pacific.
370. Li, Y.-H., T.L. Ku, G.G. Mathieu and K. Wolgemuth. 1973.
Barium in the Antarctic Ocean and implications regarding themarine geochemistry of Ba and 226 Ra . Earth and Planet. Sci.
Lett. -V19, p. 352-358.
The relationship between the distributions of bariumand 226 Ra in the Antarctic Ocean.
371. Mann, C.R., A.R. Coote and D.Mo Gardner. 1973. The meridional distribution of silicate in the western Atlantic Ocean.
Deep Sea Res. V20, p. 791-802.
Demonstration that the distribution of high dissolved
silicate water closely parallels that of the lowtemperature and low salinity Antarctic Bottom Water.
372. Cormack, E.C. 19730 Silicate and potential temperature in
the deep and bottom waters of the western Weddell Sea. Deep
Sea Res. V20, p. 927-932.
A discussion of potential temperature-silicate correlations as a means of distinguishing the water masses ofthe Weddell Sea.
373. Chung, Y.-C. 1974. Radium-226 and Ra-Ba relationships in
Antarctic and Pacific Waters. Earth & Planet. Sci. Lett. V23,
p. 125-135.
Relationship between the distributions of barium and625

226 Ra in the Antarctic Ocean.
374. Chan, L.H., J.M. Edmond, R.F. Stallard, W.S. Broecker, Y.-C.
Chung, R.F. Weiss and T.L. Ku. 1976. Radium and barium at
GEOSECB stations in the Atlantic and Pacific. Earth & Planet.
Sci. Lett. V32, p. 258-267.
C.omparison of the distributions of radium and bariumin the sea.
375. Chan, L.H., D. Drummond, J.M. Edmond and B. Grant. 1977. On
the barium data from the Atlantic GEOSECS Expedition. Deep
Sea Res. V24, p. 613-649.
Presentation and discussion of the distribution of
barium in the Atlantic Ocean based on measurements
made on samples collected as part of the Atlantic
GEOSECS expedition.
376. Edmond, J.M., S.S. Jacobs, A.L. Gordon, A.W. Mantyla and R.F.
Weiss. 1979. Water column anomalies in dissolved silica overopaline pelagic sediments and the origin of the deep silica
maximum. Jour. of Geophys. Res. V84, p. 7809-7826.
Distribution of silicate throughout the world ocean.
377. Chung, Y. 1980. Radium-barium-silica correlations and a twodimensional radium model for the world ocean. Earth &Planet. Sci. Lett. V49, p. 319-328.
C~rrelation between the distributions of barium andsilicate in the oceans.
378. Froelich, P.N. and M.O. Andreae. 1981. The marine geochemistry of germanium: ekasilicon. Science V213, p. 205-207.
A discussion of marine geochemistry of germanium. The-geochemical behavior of dissolved germanium in rivers,-estuaries and the oceans is similar to that of silicon.

3. PAPERS DEALING WITH THE DISTRIBUTIONS OF ECD2,
C03=, C02 AND ALKALINITY IN THE SEA (#379 to '409)
379. Buch, K., H.W. Harvey and S. Grifenberg. 1932. Uber der
Kohlen - sourcesystem in Meerwasser, Rapp. Proc. Verb.
C. P . I . E. M. V79, p. 1-70 .
Discussion of the C02, HC03-, C03= equilibria in sea·water.
380. Euch, K. 1933. Der Borsauregehalt der Meerwassers and seine
Bedentung bei der Berechnung des Kohlensauresystem in Meerwasser, Conseil Perm. Intern. l'Explor. Mer. Rapp. Proc-Verb.
Ve5, p. 71-75.
The importance of boric acid to the CO 2 , HC0 3 - , C03=
equilibria in sea water.
381. Buch, K. 1933. On boric acid in sea and its influence on the
carbonic acid equilibrium. Conseil Perm. Intern. l'Explor.
Mer. Rapp. Proc-Verb. V85, p. 71-75.

626

The first paper showing the importance of taking into
account borate when considering the relationshipbetween the concentrations of CO 2 , HC0 3 -, C0 3 = andalkalinity in sea water.
382. Wattenberg, H. 1933. Uber die Titrationsalkalinitat und den
Kalziumkarbon algehalt des Meerwassers, Duetche Atlantische
Exped., Meteor 1925-1927, V8, p. 122-231.
Measurements of alkalinity and EC02 in waters from theSouth Atlantic.
383. Buch, K. 1938. New determination of the second dissociation
constant of carbonic acid in seawater. Conseil Perm. Acta
Acad. Aboenis Math. Pys. VII, p. 18.
An early but surprisingly accurate measurement of thesecond dissociation constant of carbonic acid (i.e.,K' 2) in seawater.
384. Buch, K. 1939. Beobachtungen uber dos Kohlenosaure
gleichgewicht und uber den Lohlensaursaustansch zwischen
atmosphere und Meer in nord atlantischen Ozean, Acta Acad.
Aboenis Math Phys. VII, p.5-16.
Measurements of pH, alkalinity, EC0 2, pC02 and dissolved
oxygen on surface waters in the North Atlantic over theperiod June through September 1935.
385. Buch, K. 1951. Das Kohlensaure Gleichgewichtssystem in
Meerwasser.Hansforskning Inst. Skr. No. 151, p. 18.
C02, HC03-, C03= equilibria in sea water.
386. Lyman, J. 1956. Buffer mechanism of sea water, Ph.D.
thesis, University of California, Los Angeles, p. 1-196.
-Measurements of the first and second dissociation
constants of carbonic acid in sea water and of the
dissociation constant of boric acid in sea water.
387. Takahashi, T. 1961. Carbon dioxide in the atmosphere and in
Atlantic Ocean water. J. Geophys. Res. V66, p. 477-494.
Traverses of C02 partial pressure in the atmosphere andin the surface ocean along a north-south traverse in the
Atlantic Ocean carried out as part of the International
Geophysical Year Program in 1958.
388. Keeling, C.D., N.W. Rakestraw and L.S. Waterman. 1965.
Carbon dioxide in surface waters of the Pacific Ocean 1,
measurements of the distribution. J. Geophys. Res. V70,
p. 6087-6097.
Summary of C02 partial pressure measurements in thesurface waters of the Pacific Ocean made during andafter the International Geophysical Year Program.
389. Keeling, C.D. 1965. Carbon dioxide in surface waters of
the Pacific Ocean 2, calculation of the exchange with theatmosphere. J. Geophys. Res. V76, p. 6099-6102.

627

Based on the concentration of C02 in the atmosphere andocean water, an exchange rate across sea-air interface
is estimated to be 18 mol/cm 2 yr in the Pacific Ocean.
390. Disteche, A. and S. Disteche. 1967. The effect of pressure
on the dissociation of carbonic acid from measurements withbuffered glass electrode cells. J. Electrochem. Soc. Vl14,
p. 330-340.
First direct measurements (using high pressure pH
electrodes) of the dissociation constants for sea waterat conditions approximating those in deep water.
391.

C~lberson, a., D.H. Kester, and R.M. Pytkowicz.
1967. Highpressure dissociation of carbonic and boric acids in sea
water. Science V157, p.59-61.
Another study of the pressure effects on the dissociation constants of carbonic acid in seawater.

392.

C~lberson, C.H. and R.M. Pytkowicz.
1968. Effect of pressure on carbonic acid, boric acid and the pH in seawater.
Limnol. and Oceanog. V13, p. 403-417.
A study of the pressure effects on the carbonate
chemistry of sea water.

393. Keeling, C.D. and L.S. Waterman. 1968. Carbon dioxide insurface ocean waters 3. Measurements on Lusiad Expedition
1962-1963, J. Geophys. Res. V73, p. 4529-4541.
Further measurements of the C02 partial pressure insurface ocean waters.
394. Keeling, C:.D. 1968. Carbon dioxide in surface ocean waters
4. qlobal distribution. J. Geophys. Res. V73, p. 4543-4553.
~ global synthesis of CO 2 partial pressure distribution
in the surface ocean.
395. Edmond, J.M. 1970. High precision determination of titration
alkalinity and total carbon dioxide content of sea water bypotentiometric titration. Deep Sea Res. V17, P. 737-750.
The development of a closed titration cell designed toallow high precision analyses of the alkalinity andtotal dissolved inorganic carbon content of sea water
samples.
396. Takahashi, T., R.F. Weiss, C.H. Culberson, J.M. Edmond,
D.E. Hammond, a.s. Wong, Y-H Li and A.E. Bainbridge. 1970. A
carbonate chemistry profile at the 1969 GEOSECS intercalibration station in the eastern Pacific Ocean. Jour. of Geophys.
Res. V75, p. 764e-7666.
Intercomparison of the results of pH, titration Le02
and alkalinity, gas chromatograph LC02 and C02 partialpressure measurements on samples collected at the GEOSECS intercalibration station in the eastern North
Pacific.

628397. Craig, H. and R.F. Weiss. 1970. The GEOSECB 1969 Intercalibration Station: Introduction and hydrographic features, andtotal CO 2 - 02 relationships. J. Geophys. Res. V75 p. 7641 -

7647.

'

Profile of LC02 and 02 at the GEOSECS intercalibration
station in the eastern North Pacific.

398. Brewer, P.G., G.T. Wong, M.D. Bacon, and D. Spencer.

1975.

An oceanic calcium problem? Earth and Planet. Sci. Lett.
V26, p. 81-87.
Demonstration that the uptake of nitrate by plants insurface water leads to an increase in the alkalinity ofsurface water and that the release of nitrate duringrespiration leads to a decrease in the alkalinity of
deep water.

399. Millero, F.J. and R.A. Berner.

1972. Effect of pressure on
carbonate equilibria in seawater. Geochim. et ITosmochim.
Acta. V36, p. 92-98.
A calculation based on partial molar volumes of thepressure dependence of the solubility of calcite inseawater.

400. Merbach, C., C.H. Culberson, J.E. Hawley and R.M. Pytkowicz.
1973. Measurement of the apparent dissociation constants ofcarbonic acid in sea water at atmospheric pressure, Limnol.
and Oceanog. VIS, p. 897-907Redetermination of the dissociation constants ofcarbonic acid. These constants are currently acceptedby most workers.

401. Hausson, I. 1973. A new set of acidity constants for carbonicacid and boric acid in sea water. Deep Sea Res. V20, p.

461-478.

Yet another determination of the dissociation constantsfor carbonic acid in sea water.

402. Gieskes, J. 1974.

The alkalinity - total carbon dioxidesystem in sea water, The Sea V5 ed. by E.D. Goldberg, J.
Wiley and Sons, New York, N.Y., p. 123-156.
A review artcile on measurements of rC02 and,alkalinity in sea water.

403. Weiss, R.F.

1974. Carbon dioxide in water and seawater: The
solubility of a non-ideal gas. Marine Chern. V2, p. 203-215.
Measurements of the solubility of C02 in seawater.

404. Takahashi, T., P. Kaiteris, W.S. Broecker and A.E. Bainbridge. 1976. An evaluation of the apparent dissociation
constants of carbonic acid in sea water. Earth and Planet.
Sci. Lett. V32, p. 458-467.
Use of redundant measurements of C02, EC02 and alkalinity made as part of the GEOSECS program to evaluate thereliability of various sets of equilibrium constants.

629

405. Fiadero, M. 1980. The alkalinity of the deep Pacific. Earth& Planet Sci. Lett. V49, p. 499-505.
An analysis of the GEOSECS alkalinity measurements inthe deep Pacific Ocean.
406. Takahashi, T., W.S. Broecker, A.E. Bainbridge and R.F. Weiss.
1980. Carbonate Chemistry of the Atlantic , Pacific and Indian Oceans: The results of the GEOSECS expeditions 1972-1978.
Technical Report NO.1, CU-I-80, Lamont-Doherty GeologicalObservatory, Palisades, N.Y., p. 211.
A listing of the C02, LC02 and alkalinity measurements
made during the Atlantic, Pacific, and Indian Ocean
GEOSECS expeditions: Also given are calculated in situ
C03=' HC0 3 - and CO 2 concentrations and the extent ofsupersaturation of the water with respect to calcite
and aragonite.
407. Bradshaw, A.L. and Brewer, P.G. 1980. The titration of sea
water with strong acid: An evaluation of the GEOSECS total
carbon dioxide-alkalinity potentiometric titration procedure.
In: "Workshop in Oceanic C02 Standardization, La Jolla,
Calif. Nov. 30-Dec. 1, 1979" ed. by H.G. Ostlund and D. Dryssen, U.S. Dept. of Energy, C02 Effects Research and Assessment Program, Washington, D.C
An evaluation of the procedures used to analyze the
GEOSECS titration data and suggestions for the improvement of these procedures.
0'408. Takahashi, T., W.S. Broecker, S.R. Werner and A.E.
Bainbridge. 1980. Carbonate chemistry of the surface watersof the world oceans. In: "Isotope Marine Chemistry", Chap. 15,
ed. by E.D. Goldberg, Y. Horibe and K. Sarubashi, Uchida
Rokakuho Pub. Co., Tokyo, Japan.
-Distribution of carbonate ion concentrations,alkalinity, and LC02 in the surface waters of theworld ocean as determined by the GEOSECS program.
40~.

Takahashi, To, W.S. Broecker and A.E. Bainbridge. 1981. The
alkalinity and total carbon idoxide concentration in theworld oceans. In: "Carbon Cycle Modeling", ed. by B. Bolin,
SCOPE 16, p. 271-286.
Average alkalinity and LC02 profiles for various regions'. of the ocean.

40ga.Broecker, W.S., 'T. Takahashi, P. Quay, D. Bos, D. Chipman and
M. Stuiver, in press. Carbon dioxide and radiocarbon budgets
for the equatorial Pacific Ocean and the equatorial upwellingrate. J. Geophys. Res.
An estimate of the rate of upwelling in the equatorial
Pacific based on measurements of pC0 2, ~C02 andnutrient constituents.

4.

PAPERS DEALING WITH THE DISTRIBUTION OF NITROGEN
SPEC:IES (OTHER THAN NITRATE IN THE SEA) (#410 to #421)

630

410. Delwiche, C.C. 1956. Denitrification.

In: Inorganic nitrogenmetabolism. ed. by W. D. McElroy and B. Glass, Johns Hopkins
Press, Baltimore, Md. p. 233-256.
A general discussion of the mechanism ofdenitrification.

411. Craig, H. and L. I. Gordon. 1963. Nitrous oxide in the ocean
and marine atmosphere. Geochimica et Cosmochim. Acta, V27,
p. 949-955.
First precise measurements of nitrous oxide in theatmosphere and surface ocean.

412. Goering, J.J. 1968.

Denitrification in the oxygen minimum
layer of the eastern tropical Pacific Ocean. Deep Sea Res.
V15, p. 151-164.
Evidence for the utlization of nitrate in thermocline
waters free of 02.

413. Delwiche, C:.C:. 1970. The nitrogen cycle. Scientific American, V223, p. 136-147.
A review of the nitrogen cycle in nature.

414. Wada, E. and A. Hattori. 1971. Nitrite distribution and nitrate reduction in deep sea waters. Deep Sea Res. V19, p.
123-132.

Measurements of nitrite concentrations in Sagami and
Suruga Bays, and in the central Pacific. A discussionof nitrate reduction in deep seawater is also given.

415. C:line, J.D. and F.A. Richards. 1972. Oxygen deficient conditions and nitrate reduction in the eastern tropical North
Pacific Ocean. Limnology and Oceanog. V7, p. 885-900.
~A study of nitrate utilization in the 02 free water of
the thermocline in the eastern North Pacific.

416.

J.D. 1973. Denitrification and isotopic fractionationin two contrasting marine environments; the eastern tropicalNorth Pacific and the C~rioco Trench. Ph.D. thesis, University of California, Los Angeles, 270 pp.
Nitrogen isotope fractionation accompanying the utilization of nitrate in 02 free waters.

C~ine,

417. Pierotti, D. and R.A. Rasmussen. 1976. CDmbustion as a sourceof nitrous oxide in the atmosphere. Geophys. Res. Lett. V3,
p. 265-267.
Anthropogenic production of nitrous oxide.

418. Weiss, R.F. and H. Craig. 1976. Production of atmospheric nitrous oxide by combustion. Geophys. Res. Lett. V3, p. 751753.
Anthropogenic production of nitrous oxide.

419. Yoshinari, T. 1976. Nitrous oxide in the sea.
stry V4, P. 189-202.

631

Marine Chemi-

The distribution of nitrous oxide with depth in the
North Atlantic Ocean and its anti-correlation with
dissolved oxygen.

420. Cohen, Y. and L. I. Gordon. 1978. Nitrous oxide in the oxygen
minimum of of the eastern tropical North Pacific: evidence
for its consumption during denitrification and possible
mechanisms for its production. Deep Sea Res. V-25, p. 509-

524.
The consumption of nitrous oxide under 02 freeconditions.

421. Cohen, Y. and L.I. Gordon. 1979. Nitrous oxide production in
the ocean. Jour. of Geophys. Res. V84, p. 347-353.
The production of nitrous oxide during respiration.

5. PAPERS BASED ON THE GEOSECS DATA SET DEALING WITH
THE RELATIONSHIP BETWEEN POTENTIAL TEMPERATURE,
SALINITY AND THE DISTRm UTIONS OF SEVERAL OF THE
NUTRIENT CDNSTITUENTS (1422 to '433)

422. C:raig, H., Y. Chung and M. Fiadero. 1972. A benthic front in
the South Pacific. Earth & Planet. Sci. Lett. V16, p. 50.
The use of hydrographic and nutrient element data todefine a major deep water front in the South Pacific
Ocean.

423. Broecker, W.S., T. Takahashi and Y.H. Li. 1976. Hydrographyof the central Atlantic-I: The two-degree discontinuity.
Deep Sea Res. V23, p. 10~3-1104.
The use of hydrographic and nutrient constituent data
to define a major discontinuity present throughout the-_ deep Atlantic Ocean.

424. Bainbridge, A.E. 1976. GEOSECB Atlantic Final Hydrographic

Data Report. GEOSECS Operations Group, Scripps Institution of
Oceanography, La Jolla, California.
Listings of potential temperature, salinity, silicate,nitrate, phosphate and dissolved oxygen data for the
GEOSECS Atlantic Ocean Expedition.

425. Bainbridge, A.E. 1976. GEOSECS Pacific Final Hydrographic
Data Report. GEOSECS Operations Group, Scripps Institution of
Oceanography, La Jolla, CAlifornia.
Listings of potential temperature, salinity, density,
silicate, nitrate, phosphate and dissolved oxygen data
for the GEOSECB Pacific Ocean Expedition.

426. Bainbridge, A.E. 1978.

Hydrographic Report for GEOSECS
Indian Ocean Expedition, Scripps Institution of Oceanography,
La Jolla, California.
Listings of potential temperature, salinity, silicate,nitrate, phosphate and dissolved oxygen data for the
GEOSECS Indian Ocean Expedition.

632

427. Broecker, W.S., T. Takahashi and M. Stuiver. 1980. Hydrography of the central Atlantic-II: Waters beneath thetwo-degree discontinuity. Deep-Sea Res. V27A, p. 397-419.
The relationship between nutrient constituent concentrations and salinity for waters beneath the TwoDegree Discontinuity in the western basin of the deep
Atlantic.
428. Weiss, R.F., H.G. Ostlund and H. Craig. 1979. GeochemicalStudies of the Weddell Sea. Deep Sea Res. V26, p. 1093-1120.
The use of potential temperature, salinity, isotopic
ratios, and nutrient constituent relationships todefine the contributions of various water types to theformation of Weddell Sea bottom water.
429. Broecker, W.S. and A. Bainbridge. 1978. An Abyssal shear
zone. J. of Geophys. Res. V83, p. 1963-1966.
Evidence for a thin layer of young deep water moving
along the bottom into the western basin of the North
Atlantic.
430. Broecker, W.S. and T. Takahashi. 1980. Hydrography of thecentral Atlantic-III: The North Atlantic deep-water complex.
Deep Sea Res. V27A, p. 591-613.
Nutrient constituent relationships within the North
Atlantic Deep Water complex of the Atlantic Ocean.
431. Broecker, W.S., J.R. Toggweiler and T. Takahashi. 1980. The
Bay of Bengal - a major nutrient source for the deep Indian
Ocean. Earth and Planet. Sci. Lett. V49 p. 506-512.
Demonstration that the rain of organic matter in the
Bay of Bengal constitutes a major sink for dissolved-oxygen and source for nutrient constituents in the deep
waters of the western Indian Ocean.
432. Broecker, W.S. and T. Takahashi. 1981. Hydrography of thecentral Atlantic-IV: Intermediate waters of Antarctic origin.
Deep Sea Res. V28, p. 177-193.
Demonstration that O2 minimum waters from the Antarctic Ocean penetrate along isopycnal horizons far
into the Atlantic Ocean carrying with them a strong
nutrient constituent signature.
433. Williams R.T. 1981. Hawaii-Tahiti Shuttle Experiment Hydrographic Reports Vol.I, II and III. ,Scripps Institution of
Oceanography, PAQO~ Publ.#209, 210, 211 La Jolla, Calif.
Listings of potential temperature, salinity, density,
silicate, nitrate, phosphate, dissolved oxygen, ECD2and alkalinity data taken over the period February 1979
to February 1980 along traverses from Hawaii to Tahiti.

B. CHEMICAL TRACERS OF PARTICLE-WATER INTERAC~IONSIN THE SEA (#434 to 1456)

633

OC~URRING1. PAPERS DEALING WITH THE DISTRIBUTION OF LEAD IN THE
SEA (6434 to #436)
434. Tatsumoto, M. and C.C. Patterson. 1963. The concentrations ofcommon lead in sea water. In:
Earth Science and Meteorites,
ed. by J. Geiss and E.D. Goldberg, North-Holland, Amsterdam,p. 74-89.
First reliable measurements of lead in sea water.
435. Patterson, C. 1974. Lead in seawater. Science, V183, p. 553.
Discussion of the importance of the contribution oflead from gasoline to the lead content of surface
water.
436. Schaule, B.K. and C.C. Patterson. 1981. Lead concentrationsin the northeast Pacific:
evidence for global anthropogenic
perturbations. Earth & Planet. Sci. Lett. V54, p. 97-116.
Vertical profiles of dissolved lead in the Pacific
Ocean and estimates of the relative contributions of
natural and anthropogenic lead.

2. PAPERS DEALING WITH THE DISTRIBUTION OF CADMIUM IN
THE SEA (6437 to 1438)
437. Boyle, E., F. Sclater and J.M. Edmond. 1976. On the marine
geochemistry of cadmium. Nature V263, p. 42-44.
First reliable measurements of cadmium in sea water.
Profiles showing a depletion in surface water.
438. Bruland, K.W., G.A. Knauer and J.H. Martin. 1978. Cadmium innortheast Pacific waters. Limnology and Oceanog. V23, p.
618-625.
-Highly precise measurements of cadmium in waters from-the North Pacific demonstrating the near perfect correlation between the distribution of this metal andthat of P0 4 in the sea.

3. PAPERS DEALING WITH THE DISTRIBUTION OF CDPPER IN
THE SEA (#439 to #441)
439. Boyle, E. and J.M. Edmond. 1975. C~pper in surface waters
south of New Zealand.
Nature, V253, p. 107-109.
~Demonstration that the concentration of copper insurface water increases toward the Antarctic as doesthat for the nutrient constituents.
440. Boyle, E., F.R. Sclater, and J.M. Edmond. 1977. The distribution of dissolved copper in the Pacific. Earth & Planet. Sci.
Lett. V37, p. 38-54.
Vertical profiles of copper in the Pacific Ocean showing a continuous increase from surface to sea floor.
441. Moore, R.M. 1978. The distribution of dissolved copper inthe eastern Atlantic Ocean. Earth & Planet. Sci. Lett. V41,
p. 461-468.

634

A discussion of dissolved copper concentrations at fivestations in the eastern Atlantic ,between 20° and 24°N.
4. PAPERS DEAL1NG WITH THE DISTRffi UTION OF ZINC: AND
NICKELIN THE SEA (1442 to '443)

442. Sclater, F.R., E. Boyle and J.M. Edmond. 1976. On the marine
geochemistry of nickel. Earth & Planet. Sci. Lett. V31, p.
119-128.
First reliable vertical profiles of nickel in the sea.
The concentration decreases up the water column but
does not go to zero in surface water.
443. Bruland, K.W., G.A. Knauer and J.H. Martin. 1978. Zinc innortheast Pacific water. Nature, V27l, p. 741-743.
Vertical profile of zinc in the northeast Pacific Ocean
and demonstration of its excellent correlation withthat of silica.

5. PAPERS DEALING WITH THE DISTRffiUTION OF MANGANESE IN
SEA WATER (6444 to 1449)
444. Bender, M.L., T.L. Ku and W.S. Broecker. 1970. Accumulation
rates of manganese in pelagic sediments and nodules. Earth &Planet. Sci. Lett. VB, p. 143-148.
A comparison of the rate of accumulation in Mn nodules
with that in sediments.
445. Bender, M.L., G.P. Klinkhammer and D.W. Spencer. 1977. Manganese in sea water and the marine manganese balance. Deep
Sea Res. V24, p.799-812.
First reliable measurements of the concentration of- manganese in sea water.
446. Klinkhammer, G.P., M. Bender and R.F. Weiss. 1977. Hydrothermal manganese in the Galapagos Rift. Nature, V269, p.
319-320.
Demonstration that manganese accompanies the excess 3 He
in the hot waters issuing forth from the ridge crestsin sufficient quantities to serve as a tracer for thesewaters.
44'( .

R.F. 1977. Hydrothermal manganese in the deep sea:scavenging residence time and Mn/ 3 He relationships. Earth &Planet. Sci. Lett. V37, p. 257-262.
Based on the contrast between the Mn/ 3 He ratio in ridge
crest hydrothermal waters and in ambient 3 He contaminated deep water it is shown that the residence time of
Mn in deep water can be no more than a few decades.

Weiss~

448. Klinkhammer, G.P. and M.L. Bender. 1980. The distributionof manganese in the Pacific Ocean. Earth and Planet. Sci.
Lett. V46, p. 361-384.
Profiles of manganese in the Pacific Ocean.

635449. Landing, W.M. and K. W. Bruland. 1980. Manganese in the North
Pacific. Earth and Planet. Sci. Lett. V49, p. 45-56.
Demonstration that the concentration of manganese insurface waters decreases away from the coasts in keeping with the short residence time of this element in
sea water.

6. PAPERS DEALING WITH THE DISTRIBUTION OF SELENIUM IN
SEA WATER ('450 to '451)
450. Measures, C.l. and J.D. Burton. 1980. The vertical distribution and oxidation states of dissolved selenium in the northeast Atlantic Ocean and their relationship to biological processes. Earth & Planet. Sci. Lett. V46, p. 385-396.
Profiles of selenium in the Atlantic Ocean. Also ademonstration that the two oxidation states of selenium
are not at thermodynamic equilibrium with one another.
Rather they act more like the pair N2 -N0 3 in the sea(i.e., slowly interacting.)

451. Measures, C.l., R.E. McDuff and J.M. Edmond. 1980.

Selenium
redox chemistry at GEOSECS I re-occupation. Earth & Planet.
Sci. Lett. V49, p. 102-108.
Further results on the vertical distribution ofselenium and its oxidation states.

7. PAPERS DEALING WITH SEVERAL METALS (#452 to '456)
452. Martin, J.H. and G.A. Knauer. 1973.

The elemental composition of plankton. Geochim. C.osmochim. Acta, V37, p. 1639-

1653.

The elemental composition of phytoplankton samples_collected in Monterey Bay. CAlif., is reported in thispaper.

453. Knauer, G.A. and J.H. Martin. 1973. Seasonal variations of
cadmium, copper, manganese, lead and zinc in water and phytoplankton in Monterey Bay, California. Limnol. Oceanog. V18,
p. 597-604.
Concentrations of a number of metals in waters andplants in a near-shore environment.

454. Bender~ M.L. and C. Gagner. 1976. Dissolved copper, nickel
and cadmium in the Sargasso Sea. J. Marine Res. V34, p.
327-339.

Discussion of dissolved CU, Ni, and Cd in surface anddeep waters of the Sargasso Sea.

455. Bruland, K.W. 1980. Oceanographic distributions of cadmium,

zinc, nickel and copper in the North Pacific. Earth & Planet.
Sci. Lett. V47, p. 176-198.
Contrast of the profiles of four metals in the North
Pacific Ocean.

636456. Bruland, K.W. In press. Trace elements in the sea. In: Chemical Oceanography, V8, ed. by J.P. Riley and G. Skirrow,
Academic Press, London.
Review paper on trace metal distributions in sea water.
456a.Balistrieri, L., P.G. Brewer and J.W. Murray. 1981. Scavenging residence times of trace metals and surface chemistry ofsinking particles in the deep ocean. Deep Sea Res. V28A, p.
101-121.
Comparison of field measurements with predictions basedon adsorption properties of metals as measured in thelaboratory.

IV.

PAPERS DEALING WITH BIOGENIC MATERIAL (1457 to '543)
A. PAPERS DEALING WITH BIOGENIC MATERIAL IN THE SEDIMENTS(#457 to '531)
1. PAPERS DEALING WITH THE SOLtBILITY OF CALCITE AND
ARAGONITE (1457 to #468)
457. McIntyre, W.G. 1965. The temperature variation of the solubility product of calcium carbonate in seawater. Fisheries
Res. Bd. of C~nada, Manuscript Report Series V200, pl-153.
The first "modern" measurement of the solubilities ofcalcite and aragonite in sea water.
457a.Pytkowicz, R.M., and G.A. Fowler. 1967. SolUbility offoraminifera in seawater at high pressure. Geochem. J. VI,
p. 169-1t32.
Discussion of the effect of pressure on carbonate ion
pairs, on the relative solubilities of calcite andaragonite, and on rates of solution of these phases.
458. Hawley, J.E. and R.M. Pytkowicz. 1969. SolUbility of calcium
carbonate in seawater at high pressure and 2°C.. Geochim.
C~smochim. Acta. V33, p. 1557-1561.
The first measurement of the solUbility of calcium
carbonate in sea water at prsssures corresponding tothose in deep water.
458a.Millero, F.J., and R.A. Berner. 1972. Effect of pressure on
carbonate equilibria in seawater. Geochim. Cnsmochim. Acta.
V36, p. 92-98.
A discussion of the effect of pressure on carbonate
equilibria in seawater based on the partial molar
volumes.
458b.Duedall, I.W. 1972. The partial molal volumes of calcium
carbonate in sea water. Geochim. Cosmochim. Acta., V36, p.
729-734.
Experimental determination of the partial molal volume
K2 C0 3 and Na2C03 for sea water and for a.725M NaClsolution.

637

459. Ingle, S.E., C.H. Culberson, J. Hawley and R.M. Pytkowicz.
1973. The solubility of calcite in sea water at atmosphericpressure and 35 0 /00 salinity, Marine Chemistry VI, p.
295-307.
The first precise measurement of the solubility ofcalcite in sea water. These results have stood thetest of time.
460. Ingle, S.E. 1975. The solubility of calcite in the oceans.
Mar. Chern. V3, p. 301-319.
The apparent solubility product of calcite in sea water
was measured as a function of temperature, salinity,and pressure using potentiometric saturometry
techniques.
461. Berner, R.A. 1976. The solubility of calcite and aragonite insea water at atmospheric pressure and 34.5 0 /00 salinity.
Am. J. Sci. V216, p.113-730.
Measurements of the solubility of calcite and aragonite. The calcite value is considerably higher thanthat of Ingle.
462. Broecker, W.S. and T. Takahashi. 1971. The solubility of calcite in sea water. In: Thermodynamics in Geology, ed. by
D.G. Fraser, D. Reidel Pub. Co., p. 365-379.
A review of our knowledge of the solubility of calcite
and aragonite in sea water.
463. Broecker, W.S. and T. Takahashi. 1978. The relationshipbetween lysocline depth and in situ carbonate ion concentration. Deep Sea Res. V25, p. 65-95.
An attempt to determine the solubility of calcite and~aragonite in the deep sea using saturometermeasurements and lysocline depth observations.
464. Sayles, F.L. 1980. The solubility of CaC03 in seawater at 2°C
based upon in situ samples pore water composition. Marine
Chemistry, V9, p. 223-235.
Estimates of the apparent solubility products of sedimentary calcite and aragonite in seawater based uponmeasurements on water collected in situ from the poresof deep sea sediments. A particularly nice piece of'·research.
465. Morse, J.W., A. Mucci and F.J. Millero. 1980. The solubilityof calcite and aragonite in seawater at 35 0 /00 salinity at
25°C. and atmospheric pressure. Geochim. Cosmochim. Acta, V44,
p. 85-94.
Careful long term equilibrations of calcite and aragonite with seawater yielding solubility estimates consistent with those of Ingle.

466. Plath, D.C., K.S. Johnson and R.M. Pytkowicz.

1980. The
solUbility of calcite-probably containing magnesium-in sea

638

water.

Marine Chern. V10, p. 9-29.
More precise measurements of the solubility of calcite
at atmosphe~ic pressure as a function of temperatu~eand salinity.

467. Plath, D.C. and R.M. Pytkowicz. 1980. The solubility ofaragonite in seawate~ at 25.0 o C and 32.62%0 salinity.
Marine Chern. V10, p. 3-7.
A measurement of the solubility of aragonite which
agrees with the Berner value. It yields an aragonite
to marine calcite solubility ratio of 2.05 opposed tothat of 1.48 predicted by thermodynmamic consideration.
The suggestion is made that marine calcite has a lower
solubility than the pure calcite.
468. Plummer, L.N. and E.T. Sundquist. 1982. Total individual ion
activity coefficients of calcium and carbonate in seawater at
25°C and 35 0 /00 salinity, and implications to the agreement
between apparent and thermodynamic constants of calcite andaragonite. Geochim. et Cosmochim Acta V46, p. 247-258.
An attempt to reconcile the solubilities for calcite
and aragonite in sea water with those in fresh water.
2. PAPERS DEALING WITH CaC03 SATUROMETRY (#469 to'472)
469. Weyl, P.K. 1961. The carbonate saturometer. J. of Geology,
V69, p. 32-43.
The theory behind carbonate saturometry.
470. Weyl, P.K. 1965- The solution behavior of carbonate
materials in seawater. Proceedings of International Conference -on Tropical Oceanography, Miami Beach, Florida, p. 178228.
Demonst~ation using the carbonate saturometer that theformation coatings of impure calcite (containing Mg)
during calcite precipitation and perhaps even during
dissolution makes difficult the achievement ofthermodynamic equilibrium.
471. Ben-Yaakov, S. and I.R. Kaplan. 1971. Deep sea in situ calcium'~arbonate saturometry. J. of Geophys. Res. V76, p. 772731.
The development of an automated saturomater which canoperate at great depth in the sea.
472. Ben-Yaakov, S. E. Ruth and I.R. Kaplan. 1974. C.arbonate compensation depth: relation to carbonate solubility in ocean
waters. Science, V184, p. 982-894.
The application of in situ saturometry to determine therelationship between-the lysocline depth and calcite
saturation depth.

639

3. PAPERS DEALING WITH THE RESULTS OF CALCITE AND
ARAGONITE IN SITU DISSOLUTION EXPOSURES (#413 to #476)
473. Peterson, M.N.A. 1966. Calcite: rates of dissolution in a
vertical profile in the central Pacific. Science, V154, p.
1542-1544.
The first demonstration that the long-term exposure ofcalcite crystals to sea water on deep sea moorings
offered valuable information regarding the depth
dependence of dissolution rates and the depth of thehorizon of calcite saturation.

474. Milliman, J.D. 1975. Dissolution of aragonite, Mg-calcite andcalcite in the North Atlantic Ocean. Geology, V3, p 461-462.
A repeat of the Peterson experiment in the Atlantic
using aragonite and high Mg calcite as well as calcite.

475. Bonjo, S. and J. Erez. 1978.

Dissolution rates of calcium carbonate in the deep ocean: an in situ experiment in the North
Atlantic. Earth and Planet. Sci. Lett. V40, p. 226-234.
An improvement in the Peterson technique involving pumping of water across mineral grains. In so doing the influence of changing water turbulence with depth is eliminated.

476. Thunell, R.C., S. Bonjo and R. Keir. 1980. Calcite dissolution: an in situ study in the Panama Basin. EOS {Trans.
Amer. Geophys. Union) V61, p. 275.
Comparison of the calcite dissolution rates measured using the Honjo procedure wi th the pe rcent .cal ci te in sediments taken at various water depths in the same area.

- 4. LABORATORY STUDIES OF THE DEPENDENCE OF THE RATES OFCALCITE AND ARAGONITE DISSOLUTION RATE ON THE CARBONATE ION CDNTENT (HENCE pH) OF THE WATER (#477 to'419)
477. Morse, J.W. and R.A. Berner.

1972. Dissolution kinetics of
calcium carbonate in seawater II. A kinetic origin for the
lysocline. Amer. J. Science, V272, p. 840-851.
Measurements of the rate of dissolution of calcite
dispersed in rapidly stirred water at a variety of pH's',(and hence CD3= ion concentrations). A nonlinear
dependence is observed. Implications to calcite
dissolution in sediments are given.

478. Ku, T.-L. and T. Oba. 1978. A method for quantitative evaluation of carbonate dissolution in deep-sea sediments and its
application to paleoceanographic reconstruction. Quaternary
Res. VIa, p. 112-129.
An attempt to calibrate dissolution indices (i.e., thefragmentation of solution sensitive species) throughlaboratory studies employing tracer foraminifera.

640479. Keir, R.S. 1980. The dissolution kinetics of biogenic calcium carbonates in seawater. Geochim. et Cosmochim. Acta.,
V44, p. 241-252.
Demonstration by laboratory experiments that the rateof dissolution of calcite goes with the fourth power of
C03= ion concentration for rapidly stirred suspended
particles.

5. PAPERS DEALING WITH THE DISTRIBUTIONS OF CALCITE AND
ARAGONITE IN RECENT SEDIMENTS (#480 to 1495)
480. Chen, C. 1964. Pteropod ooze from Bermuda Pedestal. Science
V144, p.60-62.
Occurrence of aragonite in sediments shallower than 3kilometers in the North Atlantic Ocean.
481. Berger, W.H. 1967. Foraminiferal ooze: solution at depths.
Science V156, p. 383-385.
Demonstration that the shells of some species of nonspinose foraminifera are more resistant to dissolution
than those of spinose species.
482. Berger, W.H. 1968. Planktonic foraminifera: selective solution and paleoclimatic implications. Deep-Sea Res. V15, p.
31-43.
Application of the selective preservation of the shellsof various species of foraminifera to the distributionof calcite dissolution in sediments from the floor of
the Pacific Ocean.
483. Berger, W.H. 1970. Planktonic foraminifera: selective solution and the lysocline. Marine Geology V8, p. 111-138.
- Geographic distribution of dissolution patterns in the- deep Pacific.
484. Parker, F.L. and W.H. Berger. 1971. Faunal and solution patterns of planktonic foraminifera in surface sediments of theSouth Pacific. Deep Sea Res. V18, p. 73-107.
The influence of partial dissolution on the distribution of foraminifera shells in the recent sediments of
the South Pacific.
485. Valencia, M.J. 1973. Calcium carbonate and gross size analysis of surface sediments, western equatorial Pacific. Pacific
Science V27, p. 290-303.
Depth dependence of the calcite content of recent
sediments in the western equatorial Pacific.
486. Savin, S.M. and R.G. Douglas. 1973. Stable isotope and magnesium geochemistry of recent planktonic foraminifera fromthe South Pacific. Geol. Soc. Amer. Bull. V84, p. 2327-2342.
The influence of magnesium content on the susceptibility of foraminifera shells to dissolution.

641

487. Broecker, W.S. and S. Broecker. 1974. Carbonate dissolution
on the western flank of the East Pacific Rise. In: Studies inpaleooceanography, Society of Economic Paleontologists andMineralogists, Special Publication #20, ed. by W.W. Hay, p.
44-51.
Demonstration that the boundary between sediments
experiencing large-scale calcite dissolution andminimal calcite dissolution is quite sharp on the west
flank of the East Pacific Rise and that this horizon
was deeper during glacial time.
488. Melquen, M. and J. Thiede. 1974. Facies distribution and dissolution depths of surface sediment components from the Vema
Channel and the Rio Grande Rise (southwest Atlantic Ocean).
Marine Geology V17, p. 341-353.
Demonstration of a sharp boundary between sedimentsrich and poor in calcite in the western basin of theSouth Atlantic and that this boundary lies in theregion of transition from North Atlantic Deep Water andAntarctic Bottom Water.
489. Biscaye, P.E., V. Kolla and K.K. Turekian. 1976. Distributionof calcium carbonate in surface sediments of the Atlantic
Ocean. Jour. of Geophys. Res. V81, p. 2595-2603.
Contains a map showing the percent CaC03 in recent
Atlantic Ocean sediments.
490. KOlla, V., A.W.H. Be and P.E. Biscaye. 1976. Calcium carbonate distribution in the surface sediments of the Indian
Ocean. Jour. of Geophys. Res. V81, p. 2605-2616.
Contains a map showing the percent CaC03 in recent
Indian Ocean sediments.
491. Berger, W.H., C.G. Adelseck, Jr. and L.A. Mayer. 1976. Distribution of carbonate in surface sediments of the Pacific
Ocean. Jour. Geophys. Res. V81, p. 2617-2627.
Contains a map showing the percent CaC03 in recent
Pacific Ocean sediments.
492. Berner, R.A., E.K. Berner and R.S. Keer. 1976. Aragonite
dissolution on Bermuda pedestal: its depth and geochemicalsignificance. Earth and Planet. Sci. Lett. V30, p. 169.
'The depth dependence of the aragonite content of sediments on the Bermuda pedestal.
493. Berner, R.A. 1977. Sedimentation and dissolution of pteropods
in the ocean. In: The Fate of Fossil Fuel CD 2 in the Oceans,ed. by N.R. Andersen and A. Malahoff, Plenum Press, New York,
p. 243-260.
An estimate of the relative rain rates of calcite andaragonite in the world ocean.

494. Berger, W.H. 1978. Deep-sea carbonate: pteropod distribution
and the aragonite compensation depth. Deep Sea Res. V25, p.
447-452.

642

Estimates of the aragonite compensation depth invarious parts of the ocean.
495. Berner, R.A. and S. Honjo. 1981. Pelagic sedimentation ofaragonite: its geochemical significance. Science V211, p.
940-942.
A revised estimate of the ratio of aragonite to calcite
rain rate in the world ocean.

6. PAPERS DEALING WITH GLAC:IAL TO INTERGLACIAL CHANGESAND ARAGONITE DISSOLUTION (1496 to '512)
496. Arrhenius, G. 1952. Sediment cores from the East Pacific.
Swedish Deep Sea Expedition (1947-1948) Reports V5, pl-227.
Analysis of CaC03 content, foram fragment to wholeforam ratios and planktonic to benthic foram ratioswith depth in a series of equatorial Pacific cores.
497. Ruddiman, W.F. and B.C. Heezen. 1967. Differential solutionof planktonic foraminifera. Deep Sea Res. V14, p. 801-808.
Down core studies in the equatorial Atlantic Ocean
showing the change in preservation of solution sensitive foraminifera with climate.
498. Chen, C:. 1968. Pleistocene pteropods in pelagic sediments.
Nature, V219, p. 1145-1147.
Measurements showing a preservation peak in the abundance of pteropods in early post glacial time.
499. Hays, J.D., T. Saito, N.D. Opdyke and L.H. Burckle. 1969.
Pliocene Pleistocene sediments of the equatorial Pacific:
their paleomagnetic, biostratigraphic, and climatic record.
Geol. Soc. Amer. Bull. V8, p. 1481-1514.
Calcium carbonate cycles in long equatorial Pacificcores.
500. Hays, W.W. 1970. Calcium carbonate compensation. Deep Sea
Drilling Project Initial Reports V4, p. 672.
Evidence for long term changes in the depth of thecalcium carbonate compensation horizon.
501. Broecker, W.S. 1971. Calcite accumulation rates and glacial
to inter-glacial changes in oceanic mixing. In: The Late
Cenozoic Glacial Ages, Ed. by K.K. Turekian, Yale Univ. Press
New Haven, Conn. p. 239-265.
A discussion of the factors influencing of the calcium
carbonate budget in the ocean and an examination of therecord of budgetary changes kept in deep sea cores.
502. Hays, J.D., A. Perruzza. 1972. The significance of calcium
carbonate oscillations in eastern equatorial Atlantic deepsea sediments for the end of the Holocene warm interval.
Quaternary Res. V2, p. 355-3622.
Demonstration that the changing dilution of calcium
carbonate with wind-borne dust from Africa is the cause643

of the calcium carbonate cycles in cores from the eastflank of the mid-Atlantic Ridge.

503. Diester-Haass L., H.-J. Schrader and J. Thiede. 1973. Sedimentalogical and paleoclimatological investigations of two
pelagic ooze cores off C~pe Barbas, North-West Africa. Meteor
Forsch Ergeben C., V7, p. 19-66.
Evidence for an aragonite preservation spike at theclose of the last glacial period in cores along theAfrican continental margin.

504. Thompson, P.R. and T. Saito. 1974. Pacific Pleistocene sediments: planktonic foraminifera dissolution cycles and geochronology. Geology V2, p. 333-335.
Discussion of dissolution cycles in the Pacific Ocean
sediments based on the ratios of solution-susceptible
to solution-resistant planktonic foraminifera.

505. Damuth, J.E. 1975. Quaternary climate change as revealed bycalcium carbonate fluctuations in western equatorial Atlantic
sediments. Deep Sea Res. V22, p. 725-7.43.
Discussion of quaternary climate changes based on
fluctuation of total CaC03 content in deep sea cores.

506. Luz, B. and N.J. Shackleton. 1975.

CaC03 solution in thetropical east Pacific during the last 130,000 years. In: Dissolution of Deep-Sea 'carbonates, Special Publication #13,
Cnshman Foundation for Foraminiferal Research, p. 142-150.
The relationship between changes in the extent ofcalcite dissolution and changes in the 18 0 /16 0 ratio
for benthic foraminifera (i.e. with changes in ice
volume).

507. Gardner, J.V. 1975. Late Pleistocene carbonate dissolution
cycles in the eastern equatorial Atlantic. In: Dissolutionof Deep Sea carbonates, Ed. by W. V. Sliter, A.W.H. Be, andW.H. Berger. Cushman Foundation for Foraminiferal Res.
Special PUblication #13, p. 129-141.
Evidence for glacial to interglacial changes in theextent of dissolution of the calcite accumulating in
the eastern equatorial Atlantic.

508. Van Andel, T.H. 1975. Mesozoic-C.enozoic calcite compensation
depth and the global distribution of calcareous sediments.
Earth & Planet. Sci. Lett. V26, p. 187-194.
A reconstruction of the CaC03 compensation depth in
the three major oceans for the last 60 million years.

509. Hartmann, M., P.J. Muller, E. Suess and C.H. van der Weijden.
1976. Chemistry of Late Quaternary sediments and their interstitial waters from the N.W. African continental margin.
Meteor Forsch. Ergebnisse C. No. 24, p. 1-67.
Discussion of temporal changes in the lithology ofcontinental margin sediments from the eastern North
Atlantic Ocean.

644

510. Thunell, R.C. 1976. Calcium carbonate dissolution in Late

Quaternary deep sea sediments, western Gulf of Mexico.
Quaternary Res. V6, p. 281-297.
Discussion of the degree of CaC03 dissolution based on
foraminiferal test fragmentation, benthonic foram
abundance, CaC03 content, and various relationships
between solution-resistant and solution-susceptible
species in deep sea cores.

511. Berger, W.H. 1977. Deep-sea carbonate and the deglaciation
preservation spike in pteropods and foraminifera. Nature,

V269,

p.

301-303.

Evidence that the lysocline fo~ aragonite in the easttern Atlantic deepened by over 1 kilomete~ for a periodof several thousands of years corresponding to thetransition between late-glacial and post-glacial time.

512. Sundquist, E., D.K. Richardson, WeS. Broecker and T-H. Pengo
1977. Sediment mixing and carbonate dissolution in the southeast Pacific Ocean. In: The Fate of Fossil Fuel C02 in the
Oceans. Ed. by N.R. Andersen and A. Malahoff, Plenum Press,
New York p. 429-454.
An attempt to reconstruct the events which created thecore top QaC03-rich layer just beneath the present day
lysocline on the west flank of the East Pacific Rise.

7. MODELS OF THE DISSOLUTION OF CALCITE ON THE SEAFLOOR (#513 to #519)

513. Li, Y.H., T. Takahashi and W.S. Broecker. 1969. Degree ofsaturation of CaC03 in the oceans. Jour. of Geophys. Res.

V74, p. 5507-5525.
-Estimates of the degree of CaC03 saturation in the
Pacific and Atlantic Oceans based on measurements ofthe partial pressure of COz gas and the total content
of dissolved inorganic carbon content of sea water
samples.

514. Heath, GeR. and C.H. Culberson. 1970. Calcite: degree ofsaturation, rate of dissolution, and the compensation depth
in the deep oceans. Geological Soc. of Amer. Bull. V81, p.

3157-3160.

',Discussion of the hydrographic features that control
the distribution of calcite in deep-sea sediments.

515. Edmond, J.M. 1974. On the dissolution of carbonate andsilicate in the deep ocean. Deep Sea Res. V21, p. 455-480.

Discussion of the possible role of water turbulencein the dissolution of mineral phases on the sea floor.

516. Schink, D.R. and N.L. Guinasso, Jr. 1977. Modelling the influences of bioturbation and other processes on calcium carbonate dissolution at the sea floor. In: The Fate of Fossil
Fuel CD z in the Oceans, ed. by NeR. Andersen and A. Malahoff,

645

Plenum Press, New York, p. 375-400.
A model of the interaction between the rain rate of
calcite, the rain rate of other phases, the dissolution
rate of pure calcite, and the rate of bioturbation.
517. Takahashi, T. and W.S. Broecker. 1977. Mechanisms for calcite dissolution on the sea floor.
In: The Fate of Fossil
Fuel CO 2 in the Oceans, ed. by N.R. Andersen and A. Malahoff,
Plenum Press, New York, p. 455-478.
A discussion of the interaction between the potential
dissolution rate of calcite, the rain rate of calcite
and the rain rate of other phases for three different
assumptions regarding the rate limiting step for dissolution. Also a means of distinguishing among these
possible models is discussed.
518. Morse, J.W. 1978. Dissolution kinetics of calcium carbonate
in sea water: VI The near-equilibrium dissolution of calcium
carbonate rich sediments. Amer. Jour. Sci. V278, p. 344-353.
Discussion of CaC03 dissolution kinetics in CaCD 3 -rich
deep sea sediment.
519. Emerson, S. and M. Bender. 1981. Carbon fluxes at the
sediment-water interface of the deep-sea: calcium carbonate
preservation. J. Marine Res. V39, p. 139-162.
A discussion of the role of bioturbated organic material to the dissolution of calcite.
519a.Crowley, T.J., in press. Calcite dissolution patterns in thecentral North Atlantic during the last 150,000 years. Mar.
Geology.
Evidence for increases in the intensity of calcite dis- solution in the deep Atlantic during the peak glacial
episodes 70,000 and 20,000 years ago. A clear demonstration that the Atlantic and Pacific are out of phasein this regard.

8. PAPERS DEALING WITH THE DISTRIBUTION OF OPAL IN MARINE SEDIMENTS (#520 to #523)
520. Lisitzin, A.P. 1960. Bottom sediments of the eastern Antarctic and the southern Indian Ocean. Deep Sea Res. V7, p.89-99.
',Data on the opal content of sediments in the southern
Indian Ocean and adjacent Antarctic.
521. Hurd, D.C. 1973. Interaction of biogenic opal, sediment andsea water in central equatorial Pacific. Geochim. Cosmochim.
Acta V37, p. 2257-2282.
Discussion of the dissolution of biogenic opal indeep sea sediments.
522. Hurd, D.C. and F. Theyer. 1975. Changes in the Physical andChemical Properties of Biogenic Silica from the Central Equatorial Pacific I, SolUbility, Specific Surface Area, and Solution Rate Constants of Acid-Cleaned Samples. Analytical646

Methods in Oceanography, ed. by T.R.P. Gibbs, Advance in
Chemistry Series, V147, p. 211-239.
Factors influencing the solubility of opal.
522a.Nelson, D.M. and L.I. Gordon. 1982. Production and pelagic
dissolution of biogenic silica in the Southern Ocean. Geochimica and Cbsmochimica Acta V46, p. 491-501.
'Budget of dissolved silicate in waters of the Antarctic.
523. Cooke, D.W. and J.D. Hays. 1982. Estimates of Antarctic
Ocean seasonal sea ice cover during glacial intervals. In:
Antarctic Geoscience U. of Wisconsin Press, p. 1017-1025.
Glacial to interglacial changes in the distribution of
opal rich sediments in the Antarctic Ocean.

9. PAPERS CONTAINING PORE WATER PROFILES OF DISSOLVEDSILICA ('524 to '529)
524. Siever, R., K.C. Beck and R.A. Berner. 1965. Composition ofinterstitial waters of modern sediments. Jour. of Geol. V73,
p. 39-73.
Measurements of Ca, Mg, Na, K, Cl, Si02 and pH in interstitial water squeezed from samples of modern marine
sediments.
525. Fanning, K.A. and M.E.Q. Pilson. 1971. Interstitial silica
and pH in marine sediments: some effects of sampling proceduresu Science, V173, p. 1228-1231.
Discussion of temperature effects on the measurements
of interstitial silica and pH in marine sediments.
526. Fanning, K.A. and M.E.Q. Pilson. 1974. The diffusion of dissolved silica out of deep-sea sediments. J. Geophys. Res.
V79, p. 1293-1297.
Measurements of fluxes of dissolved silica from marine
sediments.
527. Heath, G.R. 1974. Dissolved silica and deep sea sediments.
In: Studies in Paleoceanography V20, ed. by W.W. Hay, Soc. of
Econ. Paleontol. and Mineral. Tulsa, Okla. p. 77-93.
A general discussion of dissolved silica in sea water'. and the silica content of deep-sea sediments.
528. Schink, D.R., N.L. Guinasso and K.A. Fanning. 1975. Processes
affecting the concentration of silica at the sediment-waterinterface of the Atlantic Ocean. J. Geophys. Res. vao,
p. 3013-3031.
Modeling pore water silica profiles.
529. Loder, T.e:., W.B. Lyons, S. Murray and H.D. McGuiness. 1978.
Silicate in anoxic pore waters and oxidation effect during
sampling. Nature, V273, p. 373-374.
Discussion of oxidation effects during sampling on thesilica concentration in anoxic pore water.

647

529a.Jahnke, R., S. Emerson, V. Grundmanis, D. Heggie and D.
Graham (in press). Pore water of the C:entral Pacific Ocean:
Nutrient Results. Earth & Planet. Sci. Lett.
Discussion of the results of the analysis of pore waterfor nutrients, ~C02, alkalinity, Mn++, Ca++ and dissolved oxygen.
529b.Sayles, F.L., F.T. Manheim and L.S. Waterman. 1973. Interstitial water studies on small core samples. Leg 15. Initial
Reports of the Deep Sea Drilling Project, V20, p. 783-804.
Listing and discussion of major constituents of pore
fluids at site 148, Leg 15 of DSDP in the C.arri bbean
Sea.
10. PAPERS CONTAINING DATA REGARDING THE DISTRIBUTION
OF ORGANIC. MATERIAL IN SEDIMENTS ('530 to 6531)
530. Bordovsky, O.K. 1965. Accumulation of organic matter inbottom sediments. Marine Geology, V3, p. 33-82.
Discussion of the accumulation of organic matter in thesediments based on information derived from the Bering
Sea.
531. Heath, G.R., T.G. Moore, Jr. and J.P. Dauphin. 1977. Organiccarbon in deep-sea sediments. In: The Fate of Fossil Fuel CD2in the Oceans, ed. by N.R. Andersen and A. Malahoff. Plenum
Press, New York, p. 605-625.
Relationship between the organic content of sediments
and their rate of deposition.

B. PAPERS DEALING WITH BIOGENIC MATERIAL IN THE WATER COLUMN('-532 to 6543)
1. PAPERS DEALING WITH SUSPENDED MATTER IN THE WATERCOLUMN (6532 to #536)
532. Brewer, P.G.,D.W. Spencer, P.E. Biscaye, A. Hanley, P.L.
Sachs, C.L. Smith, S. Kadar and J. Fredricks. 1976. The
distribution of particulate matter in the Atlantic Ocean.
Earth & Planet. Sci~ Lett. V32, p. 393-402.
Estimates of the total particle load based on the fil. tration of 10-liter water samples obtained in Niskin'bottles as part of the GEOSECS Atlantic expedition~
533. Krishnaswami, S. and M.M. Sarin. 1976. Atlantic surface particulates: composition, settling rates and dissolution in thesea. Earth and Planet. Sci. Lett. V32, p. 430-441.
Analyses of very large filters through which large
amounts of water passed during traverses from station
to station along the Atlantic GEOSECS tracks.
534. Biscaye, P.E. and S.L. Eittreim. 1977. Suspended particulateloads and transports in the nepheloid layer of the abyssal
Atlantic Ocean. Mar. Geol. V23, p. 155-172.

648

Using nephelometer calibrations obtained by filtering
water samples from Niskin bottles, the large collectionof Lamont-Doherty Geological Observatory nephelometerrecords was analyzed to yield distributions ofsuspended particle loads in the Atlantic Ocean.
535. Bishop, J.K.B., J.M. Edmond, D.R. Ketten, M.P. Bacon, andW.B. Silker. 1977. The chemistry, biology and vertical fluxof particulate matter from the upper 400 m of the equatorial
Atlantic Ocean. Deep Sea Res. V24, p. 511-548.
The analysis of samples of suspended particulate matter
obtained using an in situ pumping system coupled to alarge-diameter filtering system.
536. Dehairs, F., R. Chesselet and J. Jedwab. 1980. Discrete suspended particles of barite and the barium cycle in the open
ocean. Earth & Planet. Sci. Lett., V49, p. 528-550.
Discovery of discrete barite crystals in the material
obtained by filtering water from Niskin bottles during
the GEOSECS program.
2. PAPERS DEALING WITH MATERIAL CAUGHT IN SEDIMENTTRAPS (#537 to #541)

537. Soutar, A., S.A. Kling, A. Crill, E. Duffrin and K.W.
Bruland. 1977. Monitoring the marine environment through sedimentation. Nature, V266, p. 136-139.
One of the first papers showing the power of the sediment trap method for studies of marine chemistry andmineralogy.
538. Honjo, S. 1978. Sedimentation of material in the Sargasso Sea
at a ~5367m deep station. J. Mar. Res. V36, p. 469-492.
~The results from the first major open ocean sediment
trap deployment.
539. Brewer, P.G., Y. Nozaki, D.W. Spencer and A.P. Fleer. 1980.
Sediment trap experiments in the deep North Atlantic:
isotopic and elemental abundances. J. Marine Res. V38, p.
703-728.
Estimates of the rain rates of clay, CaCD3, opal
and organic matter based on sediment traps deployed in.. the north temperate and equatorial Atlantic Ocean.
540. Honjo, S. 1980. Material fluxes and modes of sedimentation in
the mesopelagic and bathypelagic zones. Jour. Mar. Res. V38,
p. 53-97.
Further results from sediment trap deployments.
541. Honjo, S. S.J. Manganini, J.J. CDle. 1982. Sedimentation ofbiogenic matter in the deep ocean. Deep Sea Res., V29, p.
609-625.
Chemical composition of the organic material caught insediment traps.

649

3. PAPERS DEALING WITH MODES OF TRANSPORT OF PARTICULATE MATTER T~ROUGH THE WATER CDLUMN ('542 to #543)
542. Honjo, S., and M.R. Roman. 1978. Marine copepod fecal pellets: production, preservation and sedimentation. J. Mar.
Res. V36, p. 45-57.
Studies of fecal pellets from laboratory cultures ofcommon marine copepods.
543. Honjo, S. 1980. Material fluxes and modes of sedimentation in
the mesopelagic and bathypelagic zones. J. Mar. Res. V38, p.
53-97.
Discussion of material fluxes and modes of sedimentation based on samples collected during sediment
trapping experiments at deep ocean stations.

v. PAPERS DEALING WITH CDNTROLS ON OCEAN CHEMISTRY (6544 to '547)
A. PAPERS CONCERNED WITH THERMODYNAMIC CONTROLS ON OCEANCHEMISTRY (#544 to #546)
544. Holland, H.D. 1965. The history of ocean water and its effecton the chemistry of the atmosphere. Proceedings of National
Academy of Sciences, V53, p. 1173-1182.
A general discussion of the factors controlling thechemistry of ocean water.
545. Sillen, L.G. 1966. Regulation of 02, N2 and C02 in theatmosphere: Thoughts of a laboratory chemist. Tellus V18,
p. 198-206.
Chemical equilibria between constituents in theatmosphere and solid phases in the soils, sediments,- and rocks as determinants of the 02, N2 and C02
pressures in the atmosphere.
546. Sillen, L.G. 1967. The ocean as a chemical system. Science
V156, p.1189-1197.
Suggestion that the ionic composition of sea water iscontrolled by chemical equilibria with minerals present
in ocean sediments.

B. PAPER CONCERNED WITH KINETIC CONTROLS ON OCEAN CHEMISTRY('547)
547. Broecker, W.S. 1971. A kinetic model for the chemical composition of sea water. Quaternary Res. VI, p. 188-207.
Arguments that the concentrations of many of the constituents in the sea (and in the atmosphere) are controlled by negative feedback loops that tend to drive
the loss rate of a constituent toward balance with its
rate of gain.

VI. PAPERS CONCERNED WITH MAN-MADE PERTURBATIONS IN THE EARTH'S
CARBON CYCLE ('548 to #599)

650

A. PAPERS DEALING WITH THE PRODUCTION OF FOSSIL FUEL C02, ITS
BUILD-UP IN THE ATMOSPHERE, AND ITS DISTRffiUTION WITHINTHE ATMOSPHERE ('548 to 1562)

548. Callendar, G.S. 1938. The artificial production of carbon
dioxide and its influence on temperature. Q.J. Roy. Meteorol.
Soc. V64, p.223-240.
One of the early suggestions that the build-up of C02
from the burning of coal and oil would eventually leadto a warming of the planet.
549. Callendar, G.S. 1958. On the amount of carbon dioxide in theatmosphere. Tellus, VI0, p.243-248.
Discussion of pre-industrial C02 concentration in theatmosphere.
550. Bray, J.R. 1959. An analyses of the possible recent change inatmospheric carbon dioxide concentration. Tellus VII,
p. 220-230.
Discussion of changes in atmospheric CO 2 concentration
during the recent past.
551. Keeling, C.D. 1960. The concentration and isotopic abundances of CO 2 in the atmosphere. Tellus, V12, p. 200-203.
Precise measurements of the C02 partial pressure in theatmosphere and of the 13 C/12 C r~tio in this CO 2 .
552. Bolin, B. and C.D. Keeling. 1963. Large scale atmosphericmixing as deduced from seasonal and meridional variations ofcarbon dioxide. J. Geophys. Res. V68, p. 3899-3920.
Interhemispheric mixing rates deduced from the distribution of excess CO 2 in the atmosphere (the fossil fuel~ C02 is largely produced in the northern hemisphere).
553. Bolin, B. and W. Bischof. 1970. Variations of the carbon dioxide content of the atmosphere in the northern hemisphere.
Tellus, V22, p. 431-442.
Measurements of atmospheric C02 partial pressures.
554. Bischof, W. 1971. Carbon dioxide concentration in the upper
troposphere and lower stratosphere II. Tellus, V23, p. 558561.
~Measurements of atmospheric C02 partial pressures.
555. Keeling, C.D. 1973. Industrial production of carbon dioxidefrom fossil fuel and limestone. Tellus V25, p. 174-198.
A gathering together of records of the amounts of coal,oil and natural gas burned each year and of the C02
derived therefrom. Also an assessment of the CO 2
released during the production of cement from limestone.
556. Keeling, C.D., R.B. Bacastow, A.E. Bainbridge, C.A. Ekdahl,Jr., P.R. Gunther, L.S. Waterman and J.F.S. Chin. 1976.

651

Atmospheric carbon dioxide variations at Mauna Loa Observatory, Hawaii. Tellus, V28, p. 53B-551.
Presentations of the results of more than a decade ofmonitoring of the atmospheric CO 2 partial pressure in
the northern hemisphere.
557. Keeling, C.D., J.A. Adams, Jr., C.A. Ekdahl, Jr., and P.R.
Gunther. 1976. Atmospheric carbon dioxide variations at theSouth Pole. Tellus, V28, p. 552-564.
Presentation of the results of more than a decade of
atmospheric CO 2 partial pressure monitoring in thesouthern hemisphere.
558. Machta, L., K. Hanson, and C.D. Keeling. 1977. Atmosphericcarbon dioxide and some interpretations. In: The Fate ofFossil Fuel CD 2 in the Oceans, ed. by N.R. Andersen and A.
Malahoff. Plenum Press, New York, p. 131-144.
Suggestion that the secular increase in atmosphericCD2 is modulated by changes in sea surface temperaturein the El Nino region off South America.
559. Rotty, R.M. 1977. Global carbon dioxide production fromfossil fuels and cement A.D. 1950 - A.D. 2000, In: The Fateof Fossil Fuel CD 2 in the oceans. ed. by N.R. Andersen and A.
Malahoff. Plenum Press, New York, p. 167-181.
An update of Keeling CD2 production compilation to theyear 1974 and some scenarios for the period 1974 to
2000.
560. Rotty, R.M. 1978. Atmospheric carbon dioxide: Possible consequences of future fossil fuel use. Resources and Energy(North-Holland PUblishing Co.)., Vol. 1, p. 231-249.
-- An estimate of the energy and C02 generated per unit
weight of various fossil fuels burned.
561. Rotty, R.M. and G. Marland. 1980. Constraints on fossil
fuel use. In: "Interactions of Energy and Climate," W. Bach,
J. Pankrath, and J. Williams (eds.), D. Reidel PublishingC.o., p. 191- 212 .
An estimate of the recoverable fossil fuel carbon
reserves on earth is given.
562.

Rotty~

R.M. 1981. Distribution and changes in industrial carbon dioxide production. Abstracts, World Climate Program Conconference on Analysis and Interpretation of Atmospheric C02
Data, Bern, Switzerland, p. 123-133.
A reduction in the estimates of industrial carbon
dioxide production is given according to the revisionsof the energy content of various coals by the United
Nations.

B. PAPERS DEALING WITH CHANGES IN THE LIVING BIOMASS AND SOIL
CARBON RESERVOIRS (#563 to #574)

652

563. Whittaker, R.R. and G.E. Likens. 1973. Carbon in the biota.
In:
C~arbon and the Biosphere; ed. by G.M. Woodwell and E.V.
Pecon, U.S. Atomic Energy Comm., p. 281-302.
Discussion of the net primary productivity of theworld's biosphere and the size of the world's biomass.
564. Hall, C.A.S., A. Ekdahl, and D.E. Wartenberg. 1975. A fifteen-year record of biotic metabolism in the Northern
Hemisphere. Nature, V255, p. 136-138.
Analysis of the magnitude of the seasonal changes in
the Mauna Loa record to ascertain whether any evidence
can be found for a secular change in the productiVityof the temperate northern hemisphere biomass.
Conclusion - none to be seen.
565. Whittaker, R. R. and G.E. Likens. 1975. The biosphere andman.
In: Primary Production of the Biosphere, ed. by H.
Lieth and R.R. Whittaker, Springer-Verlag, Heidelberg, p.
305-328.
Estimates of the net photosynthesis rate for various
vegetation types and of the standing biomass of each
type. Combined with areal coverage data these data
yield an estimate of the total terrestrial primary
production and for the total standing terrestrial
biomass.
566. Lemon, E. 1977. The land's response to more carbon dioxide.
In: The Fate of Fossil Fuel C02 in the Oceans, ed. by N.R.
Andersen and A. Malahoff, Plenum Press, New York p. 97-130.
Discussion of the possible effects of higher atmospheric Q02 contents on the rate of photosysnthesis, crop
yields, and the standing terrestrial biomass magnitude.
567. Schlesinger, W. H. 1977. Carbon balance in terrestrial
detritus. Annu. Rev. Ecol. Syst. V8, p. 51-81.
Discussion of the factors influencing the amounts ofcarbon stored in soils.
568. Bolin, B. 1977. Changes of land biota and their importancefor the carbon cycle. Science, V196, p. 613-615.
An attempt to assess the magnitude of the change in, living biomass caused by current forestry practices in'both developed and developing countries.
569. Stuiver, M. 1978.
Atmospheric carbon dioxide and carbon
reservoir changes. Science V199, p. 253-258.
Discussion based on tree ring 13 C data of the reductionin terrestrial carbon reservoirs since 1850 and its
effect on the atmospheric carbon dioxide concentration.
570. Woodwell, G.M., R.H. Whittaker, W.A. Reiners, G.E. Likens,
C.C. Delwiche, and D.B. Botkin. 1978. The biota and theworld carbon budget.
Science, V199, p. 141-146.
An estimate of the magnitude of the CO 2 production653

through biomass changes related to man's activities.

571. Seiler, W. and P.J. Crutzen. 1980.

Estimates of gross andnet fluxes of carbon between the biosphere and the atmospherefrom biomass burning. Climatic Change V2, p. 207-247.
An estimate of the contribution of biomass change tothe build-up of CO 2 in the air. The role of charcoal
production by forest fires is emphasized.

572. Schlesinger, W. 1982. The world carbon pool in soil organic
matter: A source of atmospheric C02? In: The Role of Terrestrial Vegetation in the Global Carbon Cycle: Measurements byremote sensing. ed. by G.M. Woodwell, John Wiley & Sons,
New York.
An assessment of the changes in the soil carbon reservoir and the resulting contribution to the changing
atmospheric CO 2 content.

573. Miller, P.C. 1981. Carbon balance in northern ecosystems andthe potential effect of carbon dioxide induced climaticchange. Carbon Dioxide Effects Research and Assessment Program, U.S. Department of Energy, p. 1-109.
An assessment of the likely changes to be expected in
the northern ecosystems.

574. Freyer, H.D., and N. Belacy. 1981. 13C/ 12C record in northernhemispheric trees during the past half millenium - anthropogenic impact and climate superpositions. Abstracts World
Climate Program Conference on Analysis and Interpretation of
Atmospheric CD2 Data, Berni Switzerland, p. 209-215.
Measurements of 13 C/ 2 C in free standing trees anddiscussions of contributions of fossil fuel input into
- atmosphere and of climatic changes to the observedchanges.

c. PAPERS DEALING WITH THE UPTAKE OF ANTHROPOGENIC C02 BY THE
SEA(#575 to '597)

575. Revelle, R. and H.E. Suess. 1957. Carbon dioxide exchange between atmosphere and ocean and the question of an increase of
atmospheric CO 2 during past decades. Tellus, V9, p. 18-27.
, First paper dealing with the uptake of fossil fuel CO 2
'by the sea. A one-box ocean model is employed.

576. Bolin, B. and E. Eriksson.

1959. Changes in the carbon
dioxide content of the atmosphere and sea due to fossil fuelcombustion. In: The Atmosphere and the Sea in Motion, Rossby
Memorial Volume, ed. by B. Bolin, Rockefeller Inst., New
York, p. 130-142.
Use of the two box ocean model to estimate the amountof fossil fuel C02 taken up by the sea.

577. Eriksson, E.

1963. Possible fluctuations in atmospheric C02
due to changes in the properties of the sea. J. Geophys. Res.

654

V68, p. 3871-3876.
Discussion of the possible effect of variations in sea
water properties on the atmospheric CD2 concentration.
578. Broecker, W.S., Y.-H. Li and T.-H. Peng, 1971. Carbon dioxide - man's unseen artifact. In: Impingement of Man in the
Oceans, ed. by D.W. Hood, Wiley - Interscience, New York,
p. 287-324.
A consideration of the reliability of two-box models in
estimating the amount of C02 taken up by the ocean.
579. Machta, L. 1972. The role of the oceans and the biosphere in
the C02 cycle. In: The Changing Chemistry of the Oceans.
Nobel Symposium 20~ed. by D. Dryssen and D. Jagner, Almquist
and Wiksell, Stockholm, p. 121-145.
The possible causes of seasonal variations inatmospheric C02 concentration is discussed. The
conclusion is that they are mainly caused by theseasonal variation of biospheric uptake and release ofC02 to the atmosphere. The role of ocean is considered
to be less important in this regard.
580. Machta, L. 1973. Prediction of C02 in the atmosphere. In:
Carbon and the Biosphere, ed. by G.M. Woodwell and L.V.
Pecon. U.S. Atomic Energy Comm., Washington, D.C., p. 51-85.
Model predictions of future and past carbon dioxide
concentration in the atmosphere.
581. Ekdahl, C.A. and C.D. Keeling. 1973. Atmospheric carbon
dioxide and radiocarbon in the natural carbon cycle: I. Quantitative deduction from records at Mauna Loa Observatory andat t~e South Pole. In: Carbon and the biosphere, ed. by G.M.
Woodwell and E.V. Pecon. U.S. Atomic Energy Comm., p. 51-85.
Discussion of airborne fraction of anthropogenic carbon dioxide. Geochemical models are used to interpret
this atmospheric C02 data.
582. Bacastow, R.B. and C.D. Keeling. 1973. Atmospheric carbon
dioxide radiocarbon in the natural carbon cycle: II. Changes
from A.D. 1700 to 2070 as deduced from a geochemical model.
In: Carbon and the Biosphere, ed. by G.M. Woodwell and L.V.
Pecon, U.S. Atomic Energy Comm., Washington, D.C., p. ~6-l35 .
. Prediction of future changes in atmosphereic C02 concentration using a nonlinear geochemical model of theinteraction of atmospheric C02 with the oceans and landbiota.
583. Keeling, C.D. 1973. The carbon dioxide cycle: reservoir
models to depict the exchange of atmospheric carbon dioxidewith the oceans and land plants. In: Chemistry of the Lower
Atmosphere, ed. by S.I. Rasool, p. 251-329.
A general discussion of reservoir models for the C02
cycle.

655

584. Oeschger, H., U. Siegenthaler, U. Schotterer and A. Gugelman. 1975. A box diffusion model to study the carbon dioxide
exchange in nature. Tellus, V27, p. 168-192.
The first paper on the one-dimensional diffusionalmodel. As this model constitutes a considerable
improvement over the two-box ocean model for problems
involving transients it has been widely adopted for estimations of the uptake of fossil fuel C02 by the sea.
585. Bacastow, R.B. 1976. Modulation of atmospheric carbon dioxideby the southern oscillation. Nature, V261, p. 116-118.
Demonstration that the secular increase in the CO 2
content of the atmosphere is modulated by somephenomena associated with the two-year periodicity
southern oscillation.
586. Bacastow, R.B. 1977. Influence of the southern oscillation on
atmospheric carbon dioxide. In: The Fate of Fossil Fuel C02
in the Oceans, ed. by N.R. Andersen and A. Malahoff, Plenum
Press, New York, p. 33-43.
More on the correlation between anomalies in the shapeof the secular increase in atmospheric CD2 content andthe southern oscillation.
587. Newell, R.E. and B.C. Weare. 1977. A relation between atmospheric carbon dioxide and Pacific sea surface temperature.
Geophys. Res. Lett. V4, p. 1-2.
Demonstration that the secular increase in atmosphericC02 content is modulated by some process associatedwith changes in the Pacific Ocean sea surface
temperature.
588. Kee11pg, C.D. and R.B. Bacastow. 1977. Impact of industrial
gases on climate. In: Energy and Climate. National Academyof Sciences, Washington, D.C., p. 72-95.
A review of predictions of future atmospheric CO 2 content rises associated with the burning of fossil fuels.
589. Zimen, K.E., P.Offermann and G. Hartmann. 1977. Source
functions of CO 2 and future C02 burden in the atmosphere. Z.
Naturforsch., V32a, p. 1544-1554.
Discussion of various source functions for C02 and"prediction of future CO 2 concentration.
590. Bolin, B.
1977. Modeling the oceans and ocean sediments
and their response to fossil fuel carbon dioxide emissions.
In: The Fate of Fossil Fuel CD2 in the Oceans. N.R. Andersen
and A. Malahoff, (Eds.), Plenum Press, New York, p. 81-95.
Toward a more detailed ocean model for the uptake offossil fuel CD2.
591. Siegenthaler, U. and H. Oeschger. 1978. Predicting futureatmospheric carbon dioxide levels. Science, V199, p. 388-395.
Use of the one-dimensional diffusion model to predict656

future atmospheric C02 contents for various energy
scenarios. Also the problem is inverted to show what
energy scenario would be required to yield a given
atmospheric C02 content trend.
592. Bacastow, R.B. 1979. Dip in the atmospheric C02 level
during the mid-1960·s. J. Geophys. Res. V84, p. 3108-3114.
Demonstration that the dip in the rise of atmosphericC02 during the mid-1960's can be explained by a coolingof the ocean surface during this period. The onedimensional diffusion model is used to quantify this
explanation.
593. Bacastow, R.B. and C.D. Keeling. 1978. Models to predict
future atmospheric C02 concentrations, In: Workshop on the
Global C~rbon Dioxide from Fossil Fuels, W.P. Elliott,& L.
Machta (Eds.), U.S. Dept. of Energy, CDNF-770385, Washington, D.C:. p. 72-90.
Prediction of future C02 concentrations.
594. Broecker, W.S., T. Takahashi, H.J. Simpson and T.-H. Pengo
1979. Fate of fossil fuel carbon dioxide and the global
carbon budget. Science, V206, p. 409-418.
A review of our knowledge of ocean uptake of fossil
fuel CD2. The conclusion is that no more than 50% of
the C02 produced to date can have taken up residence inthe ocean and therefore that the forests and soil
reservoirs have not been changing significantly overthe last 20 years.
595. Broecker, W.S., T.-H. Peng and R. Engh. 1980. Modeling thecarbon system. Radiocarbon, V22, p. 565-598.
Use of a regional one-dimensional advection-diffusion- model calibrated using the distributions of bomb
radiocarbon and tritium to estimate the uptake offossil fuel C02 by the oceans.

596. Killough, G. G. and W.R. Emanuel. 1981. A comparison ofseveral ocean models of carbon turnover in the ocean withrespect to their distributions of transit time and age, andresponses to atmospheric C02 and 14 C. Tellus, V33, p. 274290 .
.. Five ocean models are compared wi th respect to carbon
transit-time distribution, age distribution, andintegrated responses to histories of fossil C02 andbomb 14C injection.
597. Bolin, B.

19~1.

Carbon Cycle Modeling, J. Wiley, N.Y. (SCOPE

16) •A summary of contributions to the SCOPE workshop on
global carbon cycle modelling. For model calibration,a list of various carbon and carbon isotope data isgiven.

657

D. PAPERS DEALING WITH THE REAC.TION BETWEEN ANTHROPOGENIC CO 2
AND THE CaCO] IN MARINE SEDIMENTS (1598 to #599)

598. Broecker, W.S. and T. Takahashi.

1977. Neutralization offossil fuel C02 by marine calcium carbonate. In: The Fate ofFossil Fuel CD2 in the Oceans, N.R. Anderson and A. Malahoff,(Eds.), Plenum Press, New York, p. 213-241.
An attempt to estimate the amount of CaCD3 available
for dissolution by fossil fuel C02 and the pattern andrate of the attack of this CaC03 by man-made C02'599. Peng, T.-H. and W.So Broecker.

1978. Effect of sediment
miXing on the rate of calcite dissolution by fossil fuelC02' Geophys. Res. Lett. V5, p. 349-352.
Calculations showing that the rate of bioturbation in
deep sea sediments should be sufficiently rapid toprevent the build-up of an insoluble residue at the topof the bioturbated layer during the course of theattack of deep sea sediment CaC03 by fossil fuel CO 2 ,VII. PAPERS DEALING WITH THE GLACIAL TO INTERGLACIAL CHANGE IN
VARIOUS OCEAN PROPERTIES ('600 to #624)

A. PAPERS DEALING WITH THE RECONSTRUC.TION OF SEA SURFACETEMPERATURE (#600 to#610)

600. Imbrie, J. and N.G. Kipp. 1971.
A new micropaleontologicalmethod for paleoclimatology: application to a Late Pleistocene Caribbean core. In: The Late Cenozoic Glacial Ages,
K. K. Tu r ek i an, (Ed. ), Ya 1e Un i ve r sit y, New Ha ve n , p. 71-181.
The first application of multivarient statistical analysis to the problem of obtaining the temperature signal
from the data on the relative abundance of various
species of planktonic foraminifera in marine sediments.
A classic!

601. Hecht, A.D.

1973. A model for determining Pleistocene
paleotemperatures from planktonic foraminiferal assemblages.
Micropaleontology, V19, p. 69-77.
An alternate statistical model for deriving temperature
data from the relative abundance data for planktonic
foraminifera species in deep sea core material.

602. Shackleton, N.J. and N.D. Opdyke. 1973. Oxygen isotope andpaleomagnetic stratigraphy of equatorial Pacific core
V28-238: oxygen isotope temperatures and ice volumes on alOS-year time scale. Quat. Res., V3, p. 39-55.
The first detailed benthic oxygen isotope record from a
deep sea core.

603. Kipp, N.G. 1976.

New transfer function for estimating pastsea surface conditions from sea-bed distribution of planktonic foraminiferal assemblages in the north Atlantic.
Geol. Soc. Amer. Memoir, V145, p. 3-41.

658

Expansion of the data base on which the multivariant
analysis transfer function is based.
604. CLIMAP. 1976. The surface of the ice-age earth. Science,
V191, p. 1131-1137.
The first global map of sea-surface temperatures obtained using the Imbrie multi variant analysis procedure. The big surprise is the small magnitUde (-2.5 0 C)
of the average difference from present conditions.
605. Moore, T.C., Jr., L.H. Burckle, K. Geitzenauer, B. Luz, A.
Molina- Qruz, J.H. Robertson, H. Sachs, C. Sancetta, J.
Thiede, P. Thompson and C. Wenham. 1980. The reconstructionof sea surface temperatures in the Pacific Ocean of 18,000
B.P. Marine Micro., V5, p. 215-247.
Documentation of the reconstruction of sea-surface
temperatures by the multi variant analysis procedure in
the Pacific Ocean.
606. Hutson, W.H. and W.L. Press. 1980. A paleoecologicaltransfer function, FI-2, for Indian Ocean planktonic foraminifera. J. Paleontol., V54(2), p. 381-399.
A multi variant analysis transfer function for the calculation of sea-surface temperatures in the Indian
Ocean.
607. Moore, T.C.. , Jr., L.H. Burckle, K. Geitzenauer, B. Luz, A.
Molina-Cruz, J.H. Robertson, H. Sachs, C:. Sancetta, J.
Thiede, P. Thompson and C. Wenkam. 1980. The reconstruction of sea surface temperatures in the Pacific Ocean of
18,000 B.P. Marine Micro., V5, p. 215-247.
More on sea-surface temperature reconstructions byffiultivariant analyses.
608. Prell, W.L., W.H. Hutson, D.F. Williams, A.W.H. Be, K.
Geitzenauer, and B. Molfino. 1980. Surface circulation of
the Indian Ocean during the last glacial maximum, approximately 18,000 yr B.P. Quat. Res., V14, p. 309-336.
More on surface temperature reconstruction by themultivariant analysis procedure.
609. Moore) T.C.. , Jr., W.H. Hutson, N.G. Kipp, J.D. Hays, W.L.
Prell; P. Thompson and G. Boden. 1981. The biologicalrecord of the ice-age ocean. Paleogeography, PaleoclimatologYj Palaeoecology, V35, p. 357-370.
Distribution of faunal provinces (polar, subpolar,transitional, subtropical and eastern and western
tropical) in the world ocean at present and during thepeak of the last ice age.
610. CLIMAP Project Members, A. McIntyre, Leader LGM Project.
1981. Seasonal reconstructions of the earth's surface at thelast glacial maximum. Geol. Soc. Amer., Map and Chart
Series, No. 36.

659

Final synthesis of the effort to reconstruct of glacial
sea-surface temperature. New global temperature maps
are presented. The average temperature difference
between full glacial time and today turns out to be
only 1. 5°C:!
B. PAPERS DEALING WITH THE RECONSTRUCTION OF ICE VOLUME AND
HENCE OCEAN SALINITY (1611 to #617)

611. Broecker, W.S. and J. van Donk. 1970. Insolation changes,
ice volumes and the 18 0 record in deep-sea cores. Rev.
Geophys. Space Phys., V8, p. 169-198.
Discussion of the relationship between the timing ofclimatic changes and the changes in the seasonal
distribution of insolation caused by periodicities
in the earth's orbital parameters.
612. Bloom, A.L., W.S. Broecker, J.M.A. Chappell, R.K. Matthews
and K.J. Mesolella. 1974. Quaternary sea level fluctuations
on a tectonic coast: new 23IT Th /234 U data from the Heron
Peninsula, New Guinea. Quat. Res., V4, p. 185.
Estimates of the position of sea level at five times in
the past (- 102,000, ~ 82,000 - 62,000, ~ 40,000 and 28,000 years ago) based on 23b Th dating of an uplifted
coral reef sequence on New Guinea.
613. Shackleton, N. 1977. The oxygen isotope stratigraphic
record of the Late Pleistocene. Phil. Trans. R. Soc. Lond.
B, V280, p. 169-182.
Estimation of the ice volume change from the magnitudeof the 18 0/ 16 0 change in benthic foraminifera. The
sea level lowering at the last glacial maximum- obtained in this way is greater than 150 meters.
614. Fairbanks, R.G. and R.K. Matthews. 1978. The marine oxygenisotope record in Pleistocene coral, Barbados, West Indies.
Quat. Res., V10, p. 181-196.
An attempt to estimate the magnitude of sea-level lowering from information obtained from borings into thecoral reef sequence on the island of Barbados. The
amount of sea-level lowering during full glacial time
obtained in this way is at least 100 meters.
615. Peltier, W.R. and J.J. Andrews. 1976. Glacio-Isostatic
adjustment I: The forward problem. Geophys. J.R. Astr. Soc.
V46, p.605-646.
An attempt to explain the details of the global response of sea level to the melting of the ice sheets of
the last glacial period using a global response model.
One conclusion from this effort is that the lowering ofsea level in full glacial times must have been about 75
meters.
616. Broecker, W.S. 1978.

The cause of glacial to interglacial660

climatic change. In:"Evolution of Planetary Atmosphere andthe Earth". C.entre National d'Etudes Spatials(France), p. 165-177.
Demonstration that the relationship between sea level
and the 18 0 /16 0 record in benthic forams calibrated bytaking the sea levels and 18 0 /16 0 ratios for the six
times for which sea levels were established from the
New Guinea record leads to a sea level estimate of only
-65 meters for the last glacial maximum.

C~limatology of617. Denton, G.H. and T.J. Hughes. 1980. The last great ice
sheet. J. Wiley and Sons, Inc., New York.
A synthesis of the information on the position of theice margins during the peak of the last glacial period
and a flow model which yields an estimate of the volumeof the ice sheets lying within these boundaries. These
volumes yield a sea level lowering of about 160 meters
at peak glacial time.
C:. PAPERS DEALING WITH CHEMICAL CHANGES IN THE OCEAN AND
ATMOSPHERE (1618 to '624)

618. Duplessy, J.C., L. Chenguard, F. Vila. 1975. Weyl's theoryof glaciation supported by isotopic study of Norwegian core
Kll. Science, V188, p. 1208-1209.
18 0 data for cores from the Norwegian Sea showing thatduring the last glacial period this basin received its
bottom water from the open Atlantic (today it supplies
bottom water to the north Atlantic).
61~.

Streeter, S.S. and N.J. Shackleton. 1979. Paleocirculationof the deep North Atlantic: A 150,000-year record of benthic
foraminifera and 18 0 . Science, V203, p. 168-171.
- A method for determining the sense of the glacial tointerglacial change in the dissolved oxygen content of
deep sea water based on relative abundances of key
benthic forams. Also,the demonstration using thismethod that the 02 content of deep Atlantic water was
lower during glacial time than it is today.

620. Berner, W., B. Stauffer and H. Oeschger. 1979. Past
atmospheric composition and climate, gas parameters measured
on ic~ cores. Nature, V275, p. 53-55 .
. Reconstruction of atmospheric C02 content from the CO 2
to air ratio in air bubbles trapped in glacial ice.
Results from Greenland and Antarctica suggest that theC02 pressure was about 30% lower during glacial timethan during Holocene time.
621. Delmas, R.J., J-M. Ascencio and M. Legrand. 1980. Polar ice
evidence that atmospheric C02 29,000 yr BP was 50% of thepresent. Nature, V284, p. 155-157.
A separate study of the C02 to air ratio in glacial ice
leading to the same conclusion. The CO 2 content of theatmosphere was lower during glacial time.

661

622. Neftel, A., H. Oeschger, J. Swander, B. Stauffer and R.
Zumbrunn. 1982. Ice core sample measurements give atmospheric CDz content during the past 40,000 years. Nature V295 p.
220-223.
'
Measurements of COz to air ratio made by crushing one
cubic centimeter ice samples in a vacuum. The COz released is detected by a very sensitive laser technique.
Records for both Antarctica and Greenland are given.
Again 30% lower glacial values are found.
623. Broecker, W.S. 1981. Glacial to interglacial changes inocean and atmosphere chemistry. In: Climatic variations andvariability: Facts and Theories, A. Berger, editor, D. Reidel
Publishing C~., p. 109-120.
A demonstration that the atmospheric CDz partialpressure changes recorded in polar glaciers were very
likely caused by changes in the ratio of PO q to LCDZ in
the ocean.
624. Berner, W., H. Oeschger and B. Stauffer. 1980. Informationon the CD z cycle from ice core studies. Radiocarbon V22,
p.227-235.
Measurements of CD z content in ice cores from Camp
Century (Greenland) and Byrd Station (Antarctica).
624a.Broecker, W.S. 1ge2. Glacial to interglacial changes inocean chemistry. Progress in Oceanography, VII, 151-197.
Data from marine sediments supporting the lower COz
contents found for glacial time from ice core studies.
Also models showing that this change was due tovariations in either the P04/LCOz ratio in deep sea
water or to variations in the PIC ratio in organic~ res idues.

VII. PAPERS DEALING WITH RATES OF INPUT, RATES OF LOSS, AND RESIDENCE TIMES OF VARIOUS CONSTITUENTS IN THE SEA (#625 to
1670)

A. PAPERS DEALING WITH RIVER INPUTS TO THE SEA (#625 to
1633)

625. C:larke, P.W. 1924. The Data of Geochemistry. 5th edition.
U.S. Geological Survey Bulletin, V770, p. 1-841.
Contains a compilation of chemical analyses on rivers.
626. Livingstone, D.A. 1963. Chemical composition of rivers andlakes. In: The Data of Geochemistry, 6th ed., M. Fleischer,(Ed.), u:8. Geological Survey Professional Paper 440, p. 164.
C.ompilation of representati ve chemical data for thelake and river waters of the world.
627. MacKenzie, F.T. and R.M. Garrels. 1966. Chemical mass balance
between rivers and oceans. Am. J. Sci., V264, p. 507-525.

662

Discussion of removal processes for the excess dissolved constituents in the ocean required in order tomaintain the constancy of the chemical composition ofocean water.
628. Turekian, K. K. 1969. The oceans, streams and atmosphere. In:
Handbook of Geochemistry VI, K.H. Wedepohl,(Ed.), SpringerVerlag, Heidelberg, p. 297-323.
Residence time estimates for many of the constituentsof sea salt.
629. Garrels, R.M. and F.T. Mackenzie.
1971. Evolution of
Sedimentary Rocks, W.W. Norton, New York.
A treatment of the current rates of chemical andmechanical erosion, the sediment formed by theseproducts and the preservation of sediments formed inprevious geologic epochs.
630. Holland, H.D.
1978. The Chemistry of the Atmosphere andOceans, Wiley-Interscience, New York.
This book describes the processes which control thecomposition of the atmosphere and oceans, and explores
the functional relationshsips that determine thetransfer rates of the elements between the atmosphere,
the biosphere, the oceans, the crust and mantle.
631. Martin, J.-M. and M. Meybeck.
1979. Elemental mass-balanceof material carried by major world rivers. Mar. Chern., V7,
p. 173-206.
A discussion and listing of the estimates of averageriver particulate and dissolved matter composition.
632. Boyle, E.A., S.S. Huested, B. Grant and J.M. Edmond, (inpress). The chemical mass balance of the Amazon plume II:
copper, nickel and cadmium. J. Geophys. Res., V86, p. 80488066.
Estimates of the average Cu, Ni, and Cd concentrationsin Amazon River water.
632a.DeMaster, D.J. 1981. The supply and accumulation of silica
in the marine environment. Geochim. et Cosmochim. Acta, V45,
p. 1715-1732.
An account of the inputs and losses of silica to thesea concluding that 6 x 10 14 gm Si02/yr enter and leave"the sea.
This corresponds to a residence time of about
20,000 years.
633. Martin, J.-M. and M. Whitfield, (in press). The significanceof the river input of chemical elements to the ocean. Proceedings of NATO Advanced Re-search Institute on "Trace
metals in sea water", ed. by C.S. Wong, E.A. Boyle, K.W. Bruland, J.D. Buran and E.D. Goldberg, Plenum Press, New York.
A discussion of the chemical composition of dissolved
and particulate matter carried by rivers and of thequantity and compositon of material discharged to theocean system.

663

B. PAPERS DEALING WITH THE REMOVAL RATES OF CONSTITUENTS TO

DEEP SEA SEDIMENTS (6634 to 6635)

634. Froelich, P.N., M.L. Bender, N.A. Luedthe, G.R. Heath, T.
DeVries. 1982. The marine phosphorus cycle. Amer. J. Sci.
V282, p. 474~511.
A review of the residence time of phosphorus in the sea
based on river data and a new estimate based on therate of accumulation of phosphorus in marine sediments.

635 . Sa nt s ch i , P. H., D. Ad 1e r and M. Amdu r e r .

(in pre s s ) . Th efate of particles and particle-reactive trace metals incoastal waters: radioisotope studies in microcosms. Proceedings of the NATO Advanced Research Institute on Trace
Metals in Sea Water. C:.S. Wong, E.A. Boyle and K.W. Bruland,
J.D. Burton and E.D. Goldberg (Eds.), Plenum Press, N.Y.
A discussion of the transport of suspended particles
and particle-reactive trace metals from coastal
waters to sediments. The study was carried out inlarge controlled ecosystem tanks (MERL) designed tosimulate Narragansett Bay. Radio-tracer techniques
were employed.

C. PAPERS DEALING WITH INTERACTIONS BETWEEN SEA WATER AND
RIDGE CREST BASALTS (#636 to #643)

636. Bostrom, K. and M.N.A. Peterson.

1966. Precipitates fromhydrothermal exhalations on the East Pacific Rise. Econ.
Geol., V61, p. 1258 - 1265.
First demonstration of the role of ridge crest hydrothermal activity in the generation of metaliferous
sediments.

637. Bischoff, J.L. and F.W. Dickson. 1975. Seawater-basalt interactions at 200°C and 500 bars: implications for origin ofsea-floor heavy-metal deposits and regulation of sea water
Chemistry. Earth & Planet. Sci. Lett. V25, p. 385-387.
Laboratory measurements showing the kind of chemical
reactions to be expected if sea water were circulated
through basalt at high temperatures.

638. Wolery, T.J. and H. Sleep. 1976.

Hydrothermal circulationand geochemical flux at mid-ocean ridges. J. Geol. V84, p.

249-276.

An estimate of the total amount of water recirculated
through the ridge crests.

639. Bischoff, J.L. and W.E. Seyfried. 1978. Hydrothermal chemistry of sea water from 25°C to 350oC~. Amer. Jour. Sci. V278,
p. 838-860.
Further laboratory studies of the interaction
hot sea water and basalt.

between640. Mottl. M.J. and H.D. Holland. 1978. Chemical exchange during664

hydrothermal alteration of basalt by sea water I. Experimental results for major and minor components of sea water.
Geochim. CDsmochim. Acta, V42, p. 1103-1115.
Independent study of the interaction of hot sea waterwith basalt.
641. Edmond, J.M., J.B. Corliss and L.I. Gordon. 1979. Ridge
crest-hydrothermal metamorphism at the Galapagos spreading
center and reverse weathering. In: Ewing Symposium, ed. by
M. Talwani, V2, p. 383.
Discussion of chemical measurements made on hot waters
coming from ridge crest hydrothermal vents.
642. Edmond, J.M., C. Measures, R.E. McDuff, L.H. Chan, R. Collier, B. Grant, L.I. Gordon and J.B. Corliss. 1979. Ridge
Qrest hydrothermal activity and the balances of the major andminor elements in the ocean: the Galapagos data. Earth &Planet. Sci. Lett. V46, p. 1-18.
Estimates of the contribution of sea water-basalt
interactions at ridge crests to the oceanic budgets ofvarious elements.
643. Mottl, M.J. and H.D. Holland. 1979.
Chemical exchange
during hydrothermal alteration of basalt by seawater II.
Experimental results for Fe, Mn and sulphus species.
Geochim. Cosmochim. Acta, V43, p. 869-884.
Further laboratory measurements with regard to sea
water-basalt chemical interactions.

D. PAPERS DEALING WITH THE CHEMISTRY OF ATMOSPHERIC. AEROSOLSAND DUST (6644 to #652)
644. Herron, M.M., C.C. Langway, H.V. Weiss and J.H. Cragin.
1977.- Atmospheric trace metals and sulfate in the Greenland
ice sheet. Geochim. Cosmochim. Acta, V41, p. 915-920.
Measurements of Zn, pb, CD, V and sulfates in theGreenland Ice sheet.
645. Weiss, H.V., M.M. Herron and C.C. Langway. 1978. Naturalenrichment of elements in snow. Nature, V274, p. 352-353.
Demonstrates the enrichment of certain elements in snow
and discussion of the possible causes for this enrich·ment.
646. Buat-Menard, P. 1979. Influence de la retombee atmospheric
sur la chimie des metals en trace dans la matiere en suspension de l'Atlantic nord. Ph.D. thesis, University of Paris,
434pp.
Measurements of the concentration of trace metals inparticulate matter from the marine atmosphere and in
the deep ocean waters.
647. Buat-Menard, D. and R. Chesselet. 1979. Variable influenceof the atmospheric flux on the trace metal chemistry of665

oceanic suspended matter. Earth Planet. Sci. Lett., V42, p.
399-411.
Measurements of the concentration of trace metals inparticulate matter from the marine atmosphere and in
the deep ocean waters.
648. Boutron, C. 1979. Trace element content of Greenland snowsalong an eastwest transect. Geochim. Cosmochim. Acta, V43,
p. 1253-1258.
Measurements of trace element content in Greenlandsnows showing that certain elements are carried throughthe atmosphere in anomalous amounts.
649. Boutron, C. and C. Lorius. 1979. Trace metals in Antarctic
snows since 1914. Nature, V277. p. 551-554.
Discussion of the concentration of various trace metals
in east Antarctic snow.
650. Delmas R. and C. Boutron. 1980. Are the past variations of
the stratospheric sulfur burgen recorded in central Antarctic
snow and ice layers? J. Geophys. Res., V85, p. 5645-5649.
Discussion of sulfate variations in central Antarctic
snow and ice layers. Global sulfur pollution does notseem to be recorded in these samples.
651. Boutron, C. 19dO. Respective influence of global pollution
and volcanic eruptions on the past variations of the trace
metals content of Antarctic snows since 1880's. J. Geophys.
Res., V95, p. 7426-7432.
Measurements of a time sequence for trace metal
concentrations in snow samples from central east
Antarctica from 1890 to present.
652. Li, Y.-H. 1981. Geochemical cycle of elements and humanperturbation. Geochim. et Cosmochim. Acta. V45, p. 20372084.
A discussion of the pollution inputs of various
elements and their geochemical cycle in the ocean.

E. PAPERS DEALING WITH GAS SOLUBILITY IN SEA WATER (#653 to
1660)
653.

R.F. 1970. The solUbility of nitrogen, oxygen andargon in water and seawater. Deep Sea Res., V17, p. 721735.
Equations are given for calculating the solubility of
N2' 02, and Ar in sea water as a function of temperature and salinity.

Weisi~

654. Weiss, R.F. 1971. The effect of salinity on the solubility ofargon in seawater. Deep Sea Res. V18, pp. 225-230.
Special attention is given to the salt effect on thesolubility of Ar.

666

655. Weiss, R.F. 1971. Solubility of Helium and Neon in water andseawater. J. Chemical & Engineering Data, V16, No.2 p. 235241.
'
Equations are given to calculate the solubility of He
and Ne in sea water as a function of temperature andsalinity.
656. Wei~s, R.F. 1974. C~rbon dioxide in water and sea water: The
solubility of a non-ideal gas. Marine Chern., V2, p. 203-215.
An equation is given to calculate C02 solUbility in sea
water as a function of temperature and salinity.
657. Kester, Dana, R. 1975. Dissolved gases other than CO 2 . In:
Chemical Oceanography, VI, edited by J.P. Riley and G. Skirrow Academic Press, pp. 497-557.
A summary of solubility of gases in sea water, mostly
based on R.F. Weiss equations.
658. Weiss, R.F. and T.K. Kyser. 1978. Solubility of krypton inwater and seawater. J. Chern. Engineering Data, V23, p. 69-72.
An equation is given to calculate the solubility of Kr
in sea water as a function of temperature and salinity.
659. Weiss, R.F. and B.A. Price. 1980. Nitrous oxide solubility
in water and sea water. Marine Chemistry, VB, p. 347-359.
An equation is given for calculating the solubility of
N20 in sea water as a function of temperature andsalinity.
660. CRC Handbook of Chemistry and Physics, edited by R.C:. Weast,
CRC press, 1980.
Solubility of Xe and Rn in pure water is given.

F. PAPERS DEALING WITH MOLECULAR DIFFUSIVITY OF GASES IN SEAWATER (#661 to #666)
661. Rona, E. 1917. Diffussionsgrosse und Atomdurchmesser der Radiumemanation.Zeitschrift fur physikalische chemie, stochiometrie und verwandtschaftslehre, V92, p. 213-218.
The first measurement of radon molecular diffusivity
in water at 18°C:.
662. Himmelblau, D.M. 1964. Diffusion of dissolved gases inliquids. Chern. Review, V64, p. 527-550.
A review paper on the diffusivities of gases in water.
The equation for the dependence of diffusion
coefficients on temperature is given.
663. Thomas, W.J. and M.J. Adams. 1965. Measurement of the
diffusion coefficients of carbon dioxide and nitrous oxide inwater and aqueous solutions of glycerol. Trans. Faraday
Soc., V61, p. 668-673.
Measurements of the diffus i vi ty of C.o 2 and N 20 in
wa te r.

667

663a.Wise, D.L. and G. Houghton. 1968. Diffusion coefficients ofneon, krypton, xenon, carbon monoxide and nitric oxide inwater at 10-60 o c. Chern. Engr. Sci., V23, p. 1211-1216.
Measurements of the diffusion coefficients of Ne, Kr,
Xe, CO and NO in water.
664. Boerboom, A.J.H. and G. Kleyn. 1969. Diffusion coefficientsof noble gases in water, J. Chemical Physics, V50, pp.
1086-1087.
Diffusion coefficients of all noble gases (except
radon) in water are measured using the same technique.
The absolute values fallon the low side of previously
published data. However, the ratio of diffusivities
between gases should be O.K.
665. Hayduk, W. and H. Laudie. 1974. Prediction of diffusion
coefficients for nonelectro1ytes in dilute aqueous solutions,
AICHE Journal, V20, pp. 611-615.
A summary of available diffusion coefficients of gases
in water.
666. Muller, W. 1978. Die diffusions-, loslichkeits- undpermeabilitats- koeffizienten von radon-222 in Handelsublichen Kundststoffen, Master thesis, Institut fur Umweltphysik, Heidelberg.
A series of experiments designed to measure thediffusion coefficients of radon in various mediums.

G. PAPERS DEALING WITH THE EVOLUTION OF OCEAN CHEMISTRY (#667
to #670)667. Conway, E.J. 1942. Mean geochemical data in relation to
oceanic evolution. Proc. Roy. Irish Acad. V B48, p. 119-159.
Discussion of the factors controlling the chemicalcomposition of river waters and the influence of sea
salt aerosols on same.

668. Conway, E.J.

1943. The chemical evolution of the ocean.
Proc. Roy. Irish Acad. V B48, p. 161-212.
An analysis of the manner in which the chemicalcomposition of sea water might have evolved. A
thought-filled treatment done with considerable style.

669. Rubey, W.W. 1951. Geologic history of sea water. Bull.
Geol. Soc. Amer., V62, p. 1111-1147.
A classic paper on the history of the outgassing of thevolatiles, H20, HCl, H2S0 4 and CO 2 from the earth.
670. 1i, Y.-H. 1972. Geochemical mass balance among lithosphere,hydrosphere and atmosphere. Amer. J. Sci. V272, p. 119.
Modeling the geochemical mass balance for the constituents of the lithosphere, h~drosphere, and atmosphere,
using constraints from 180, 3 C and 34 S isotope data.

668

VIII.

BOOKS ON RELATED SUBJECTS ('671 to #684)

671. Sverdrup, H.U., M.W. Johnson and R.H. Fleming. 1942. The
Oceans, their physics, chemistry and biology. Prentice-Hall,Englewood Cliffs, N.J.
The classic treatise on the physics, chemistry andbiology of the oceans. An excellent compilation of thedata and ideas in this field up to World War II.
672. Libby, W.F. 1955. Radiocarbon dating, second edition, University of Chicago Press.
A book giving the principles of radiocarbon dating
written by the inventor of this method. As it was
written only five years after the discovery of 14 C innature, it treats only the method and the very early
work in this field.
673. Mero, J.L. 1964. The mineral resources in the sea. American
Elsevier, New York.
A compilation of information regarding the distribution, textures and chemistry of manganese nodules.
674. Danckwerts, P.V. 1970. Gas-Liquid Reactions. McGraw-Hill,
New York.
The ins and outs of the gas exchange process from theviewpoint of a chemical engineer.
675. Riley, J.P. and R. Chester. 1971. Introduction to Marine
Chemistry. Academic Press, New York.
The nuts and bolts of sea water chemistry.
676. Garrels, R.M. and F.T. Mackenzie. 1971. Evolution of sedimentary rocks. W.W. Norton, New York.
- A treatment of the current rates of chemical andmechanical erosion, the sediments formed by theseproducts and the preservation of sediments formed inprevious geologic epochs.
677. Goldberg, E.D. (editor) 1~74. The Sea V5, John Wiley andSons, New York.
A series of papers dealing with thermodynamics, gas
exchange, sediments, life, the impact of man and theorigins of the sea.
678. 'Broecker, W.S. 1974. Chemical Oceanography, Harcourt Brace
Jovanovich, New York, 214 pp.
The predecessor to this book.
,Riley, J.P. and G. Skirrow. 1975. Chemical Oceanography, VI,
2,3 & 4. Riley, J.P. and R. Chester. 1976., V5 & 6. 1978. V7.
Academic Press, London.
A series of articles on subjects in chemical oceanography authored by experts.
679a.Walker, J.C.G. 1977.

Evolution of the Atmosphere. MacMillan

669

Inc., New York.
Book discussing the evolution and present chemistry of
the atmosphere.
680. Holland, H.D. 1978. The Chemistry of the Atmosphere andOceans. Wiley-Interscience, New York.
A systematic treatment of material balances involved insetting the chemical composition of sea water and theatmosphere.
681. Imbrie J. and K.P. Imbrie. 1979. Ice Ages, Enslow PUblishers, Short Hills, New Jersey.
An excellent introduction to glacial cycles and theircauses.
682. Berner, A. 1980. Early diagenisis: a theoretical approach,Princeton University Press, Princeton, N.J.
A treatment of the chemical reactions and bioturbationin the upper part of the sediment column.
683. Schopf T.J.M. 1980. Paleooceanography. Harvard University
Press, 341pp.
A somewhat disjointed summary of evidence regarding
long-term changes in the operation of the oceanatmosphere system.
683a.Pearman, G.I. 1980. Carbon dioxide and climate: Australian
research. Australian Academy of Science, Canberra.
A summary of the carbon dioxide problem. Especially
useful are the papers on C02 ~rowth enhancement and on
the factors influencing the 1 C/ 12 C ratio in plants.
684.

IX.

Stumm~ W. and J.J. Morgan.
1981. Aquatic Chemistry (second
edition). Wiley-Interscience, New York.
A superb treatment of the physical chemistry of aqueous
systems.

ATLASES CONTAINING CHEMICAL DATA (#684a to #685g)

684a.Wyrtki, K. 1971. Oceanographic Atlas of the International
Indian Ocean Expedition. National Science Foundation, Washington, D.C.
--Sections of hydrographic and nutrient constituents in
the Indian Ocean.
685. Bainbridge, A.E. 1980. GEOSECS Atlantic Expedition, Volume 2,
Sections and Profiles, National Science .Foundation, Washington, D.C.
Sections of potential temperature, salinity, density,
dissolved oxygen, silicate, phosphate, nitrate,
alkalinity, and LC02 along the Atlantic Ocean tracksof the GEOSECS program. Also shown are property-property plots for individual stations and scatter plots
for the entire Atlantic Ocean.

670

685a.Bainbridge, A.E. 1981. GEOSECS Atlantic Expedition Volume 1
Hydrographic data 1972-1973, National Science Foundation,
Washington, D. C..
Listing of potential temperature, salinity, density,
dissolved oxygen, nutrient, total dissolved inorganiccarbon, alkalinity, dissolved nitrogen, dissolved argon
and radon data. Also given are CD 2 , HC03-, C0 3=, andcarbonate mineral saturation values calculated from theEC02 and alkalinity data.
685b.Broecker, W.S., D.W. Spencer, and H. Craig (in press). GEOSECS Pacific Expedition: Hydrographic Data, U.S. Government
Printing Office, Washington, D.C., Vol. 3.
Listing of potential temperature, salinity, density,
dissolved oxygen, nutrient, total dissolved inorganiccarbon, alkalinity, dissolved nitrogen, dissolved argon
and radon data. Also given are C02, HC03-, C03= andcarbonate mineral saturation values calculated from theEC0 2 and alkalinity data.
685c.Craig, H., W.S. Broecker and D.W. Spencer. 1982. GEOSECS
Pacific Expedition: Sections and Profiles, U.S. Government
Printing Office, Washington, D.C., Vol. 4.
Sections of potential temperature salinity, density,
dissolved oxygen, silicate, phosphate, nitrate,
alkalinity and EC02 along the Pacific Ocean tracks of
the GEOSECS program. Also shown are property-property
plots for individual stations and scatter plots for theentire Pacific Ocean.
685d.Spencer, D.W., W.S. Broecker, H. Craig and R.F. Weiss, (inprep.J. GEOSECS Indian Ocean Expedi tion: Hydrographic Data,
U.S. Government Printing Office, Washington, D.C., Vol. 5.
Listing of potential temperature, salinity, density,
total dissolved inorganic carbon, alkalinity, and radon
data. Also given are C02, HC03-, C03= and carbonate
mineral saturation values calculated from the EC02 andalkalinity data.
685e.Weiss, R.F., W.S. Broecker, H. Craig and D.W. Spencer, (inpress). GEOSECB Indian Ocean Expedition: Sections and Profiles, U.S. Government Printing Office, Wash., D.C., Vol. 6.
'Sections of potential temperature salinity, density,
dissolved oxygen, silicate, phosphate, nitrate,
alkalinity and ECD2 along the Indian Ocean tracks of
the GEOSECS program. Also shown are property-property
plots for individual stations and scatter plots for theentire Indian Ocean.
685f.GEOSECS Atlantic, Pacific and Indian Oceans Expeditions CTD
Data, (in prep.). U.S. Government Printing Office,Washington, D.C., Vol. 7.
Listing of the CTD data taken during GEOSECS program.

671

685g.GEOSECS Atlantic, Pacific and Indian Oceans Expeditions
Shorebased Data, (in prep.).
U.S. Government Printing
Office, Washington, D.C., Vol. 8.
Listing of the GEOSECS 14 C 3 H 226 Ra
13 C/12 C ,18 0/16 0 , D/H and barium re~ult~.
'

x. MISCELLANEOUS PAPERS AND LATE ENTRIES ('686 to #137)
6e6. Koblentz-Mishke, O.J., V.V. Volkovinsky and J.G. Kabanova.
1970.
Plankton primary production of the world ocean. In:
Scientific Exploration of the South Pacific, ed. W.S.
Wooster, National Academy of Sciences, Washington, D.C., p.
183-193.
Compilation of primary plant production data for theworld ocean.
687. Love, C.M. 1974. Eastropac Atlas V3, U.S. Government
Printing Office, Washington, D.C.
Data showing the vertical and longitudinal distributionof plant productivity in the equatorial Pacific Ocean.
688. Dueser, W.G.
1975. Reducing Environments. Chemical Oceanography, ed. by J.P. Riley and G. Skirrow 2nd edition, V3,
Adademic Press, London.
A general discussion of the development of anoxic
conditions in restricted basins and in parts of theopen ocean.
Estimates of the time scales of water
renewal and of the extent of decompositon of organic
matter in anoxic environments are also given.
689. Levitus, S. and A.H. Oort.
1977. Global analysis ofoceanographic data. Bull. Am. Meteor. Soc. V58, p. 1270-1284.
- A summary of the distribution of temperature andsalinity data for the world ocean.

6yO. Drever, J.I.

1977. Sea Water, Cycles of the Major Elements.
Benchmark Papers in Geology V45, Dowden, Hutchinson and Ross,Inc.
A set of reprints of "classic" papers discussing themechanisms controlling the chemical composition of sea
water.

6Y1. World"Energy Outlook, Exxon Background Series, Dec. 1980.
Projections of energy use, of the sources of thisenergy and of the geographic distribution of this use.
692. Goldman, J.C.
1980. Physiological processes, nutrient
availability and the concept of relative growth rate inmarine phytoplankton ecology. Primary Production in the Sea,
ed. by P.G. Falkowske, Plenum Publishing Corp. p. 179-193.
A discussion of the variability in the C:N:P ratios inplankton induced by changes in growth rate and nutrient
availability.

672693. Mankinen, E.A. and Dalrymple, G.B. 1979. Revised Geomagnetic
polarity time scale for the interval 0-5 m.y. B.P. Jour.
Geophs. Research V84, p. 615-626.
Update of the time scale for magnetic reversal over the
past five million years.
694. Emiliani, C., E.B. Kraus, and E.M. Shoemaker. 1981. Sudden
death at the end of the Mesozoic. Earth & Planet. Sci. Lett.
V55, p. 317-334.
A recap of which organisms were killed by the great
asteroid impact which marks the end of Cretaceous timeand a thermal theory for these extinctions.
695. Hahne, A., A. Volz, D.H. Ekhalt, H. Cosatto, W. Roether, W.
Weiss and B. Kromer. 1978. Pageoph., V116, p. 575-582.
Profile of chlorofluoromethanes in the Norwegian Sea
are compared with those of tritium.
695a.Druffel, E.M. 1981.
Radiocarbon in annual coral ring fromthe eastern tropical Pacific Ocean. Geophys. Res. Lett. V8,
p. 59-62.
Record of the 14 C/12 C in surface waters adjacent to
the Galapagos Islands an area of intense upwelling,based on measurements on ring-dated corals.
696. Santschi, P.H., P. Bower, U.P. Nyffeler, A. Azevedo and W.S.
Broecker (in press).
Measurements of the resistance tochemical transport posed by the deep sea benthic boundary
layer and their significance to benthic fluxes.
Limnology &Oceanography.
The use of dissolution from alabaster plates toestimate the thickness of the boundary "stagnant film"on the sea floor and in benthic chambers.
696a.Druffel, E.M., 1982. Banded Corals: Changes in Oceanic Carbon-14 Levels During the Little Ice Age. Science, V218, p.1319.
A record of the 14 C/12C ratio in the surface ocean
extending back to the 17th century based on
measurements on ring-dated corals.
697. Kusakabe, M., T.L. Ku, J. Vogel, J.R. Southon, D.E. Nelson
and G.' Richards, in press. Beryllium-10 profiles in seawater.
Nature.
Measurements of the lOBe concentration in sea water by
the accelerator technique.
698. Ku, T.L., M. Kusakabe, D.E. Nelson, J.R. Southon, R.G.
Korteling, J. Vogel and I Nowikow, in press. Constancy ofoceanic deposition of lOBe as recorded in manganese crusts.
Na ture.

673

Measurements of the lOBe/ 9Be ratio as a function of
depth in manganese crusts with long growth histories.
699. Siegenthaler, U., in press. Uptake of excess C02 by an outcrop-diffusion model of the ocean. J. Geophys. Res.
A one and a half dimension model of the ocean designed
to assess the enhancement of fossil fuel CO 2 uptake by
the ocean through descent along outcropping isopycnal
horizons.
700. Hogg, N.G., P. Biscaye, W. Gardner and W.J. Schmitz, Jr.
1982. On the transport and modification of Antarctic Bottom
Water in the Vema Channel. J. Mar. Res. V40, p. 231-263.
Current-meter based measurements of the flux of southern component water into the Atlantic Ocean.
701. Worthington, L.V. 1976. On the North Atlantic Circulation.
The Johns Hopkins University Press, Baltimore.
Summary of evidence for the ventilation and circulation
rates at various depths in the North Atlantic Ocean.
702. Piepgras, D.J. and G.J. Wasserburg. 1982. Isotopic Composition of Neodymium in Waters from the Drake Passage. Science,
V217, pc 207-214.
Data showing the distribution of Sm, Nd and 143 Nd /144 Nd
ratio in the world ocean. Particular attention is
given to the significance of the isotope results waterfrom the Drake Passage as they provide a means ofestimating the relative contributions of water ofAtlantic and Pacific origin to the deep circumpolar
current in the Antarctic Ocean.
703. Bryden, H.L. and R.D. Pillsbury. 1977. Variability of Deep
Flow in the Drake Passage from Year-Long Current Measurements. Jour. Phys. Oceanography, V7, p. 803-810.
Estimate of the flux of water through the Drake
Passage based on current-meter observations.
704. Nowlin, W.D., T. Whitworth, III and R.D. Pillsbury. 1977.
Structure and Transport of the Antarctic Circumpolar Current at Drake Passage from Short-Term Measurements. Jour.
of Phys. Oceanography, V7, p. 788-802.
'Estimate of the flux of water through the Drake
Passage based on current-meter observations.
705. Turner, J.V. 1982. Kinetic fractionation of carbon-13 during
calcium carbonate precipitation. Geochimica et Cosmochim.
Act a ., V46, p . 118 3- 11 91 .
Estimate of the equilibrium and kinetic fractionation
factors (at 25°C) for carbon isotopes during the formation of calcite and of aragonite from HC03- ion inaqueous solution.

674

706. Emerson, S. and M. Bender.

1981. Carbon fluxes at the sediment-water interface of the deep sea: calcium carbonate preservation. Jour. of Marine Res. V39, p. 139-162.
Demonstration that the bioturbation of organic residues
into the upper sediment column coupled with oxidationof this buried debris is an important contributor to
the dissolution of calcite in marine sediments. The
inclusion of this process alters the relationshipbetween the lysocline and saturation horizon depths.

707 · Anderson, R.F., M.P. Bacon and PeG. Brewer, in press.
Removal of 230 Th and 231 pa at ocean margins. Earth &Planet. Sci. Lett.
Evidence for the preferential removal of 231 Pa(relative to 230Th) in areas of high particle rain
rate.

708. Anderson, R.F., M.P. Bacon and P.G. Brewer, in press.

Removal of 230 Th and 231 pa from the open ocean. Earth &Planet. Sci. Lett.
Evidence for the preferential removal of 230Th(relative to 231 pa ) in areas of low particulate rain
rate.

709. Boyle, E.A., S.S. Huested and S.P. Jones. 1981. On the distribution of copper, nickel, and cadmium in surface waters of
the North Atlantic and North Pacific Oceans. Jour. ofGeophys. Res., V86, p. 8048-8066.
Distribution of copper, cadmium and nickel in thesurface waters of open ocean, coastal and upwellingareas.

710. Boyle, E.A. 1981. Cadmium, zinc, copper and barium in foraminifera tests. Earth & Planet. Sci. Lett., V53, p. 11-35.
The first demonstration that the concentrations of Zn
and Cd in the shells of planktonic foraminifera might
be used to determine glacial to interglacial changes in
the concentrations of these metals in the sea.

711. Boyle, E.A. and L.D. Keigwin. 1982. Deep circulation of theNorth Atlantic over the last 150,000 years; geochemical
evidence. Science. V217.

Records for both the 13 C/12 C and Cd/Ca variations inNorth Atlantic Deep Water (NADW) through the course of
a complete glacial cycle based on the analyses of benthic forams. Evidence is presented that the contribution of NADW relative to southern water types diminished during the glacial maxima ~70,000 and ~20,OOO years
ago.

712. Hester, K. and E. Boyle.

1982. Water chemistry control of
the cadmium content of some recent benthic foraminifera.
Nature, V298, p. 260-262.

675

Analyses of recent benthic foraminifera from various
places on the sea floor showing a linearity between the
Cd/Ca of the shells and the phosphate content of thelocal bottom water.
713. Graham, D.W., B.H. Corliss, M.L. Bender, LoD. Keigwin, Jr.
1981. Carbon and oxygen isotopic disequilibria of recent
deep sea benthic foraminifera. Marine Micropaleontology, V6,
p. 483-497.
Demonstration that no tem~oral changes have occurred in
the species to species 0 1 0 and 013 C differences forfour key benthi~ organisms. Also the relationship between AOD and 0 3 C for the species P. wuellerstorfi isshown to be close to that expected for respiration.
714. Bender, M.L., R.B. Lorens and D.F. Williams. 1975. Sodium
magnesium and strontium in the tests of planktonic
foraminifera. Marine Micropaleontology, V21, p. 448-459.
An attempt to reconstruct the temporal variations in
the Na/Ca, Mg/Ca and Sr/Ca ratio, in sea water based on
the analysis of planktonic foraminifera.
715. Graham, D.W., M.L. Bender, D.F. Williams, L.D. Keigwin, Jr.
1982.. Strontium-calcium ratios in C:enozoic planktonic
foraminifera. Geochim. CDsmochim. Acta V46, p. 1281-1292.
Evidence for small (i.e., ±15%) changes in the Sr/C~
ratio in sea water over the last 60 million years
based on the analysis of planktonic foraminifera.
716. Roether, W. and B. Kromer. 1978. Field determination of airsea gas exchange by continuous measurement of radon-222.
Pag e oph , V116 , p . 47 6- 485 .
-Description of system for radon profiling through theoceanic mixed layer.
717. Fairbanks, R.G. 1982. The origin of continental shelf andslope water in th~ New York Bight and Gulf of Maine:
evidence from H2 1 O/H 2 16 0 ratio measurements. Jour. Geophys.
Res. V87, p. 5796-5808.
18
The clear demonstration of the utility 0 O-salinity
diagrams in differentiating different sources of the..low salini ty waters found in coastal regions.
718. Streeter, S.S., P.E. Belanger, T.B. Kellogg and J.rr.
Duplessy. 1982. Late Pleistocene paleo-oceanography of theNorwegian-Greenland Sea: benthic foraminifera evidence.
Quarternary Research, V-18, p. 72-90.
Evidence in support of the idea that the Norwegian Sea
was not a source for bottom water during most of
glacial time.

676

719. Sharma, P. and B.L.K. Somayajulu. 1982.
lD:se dating of large
manganese nodules from world oceans. Earth & Planet. Sci.
Lett. V59, p. 235-244.
Growth rates of Mn nodules based on 1 ~e/ ~e ratios
confirm that the long term growth rate of these objects
1s generally in the range 1 to 4 mm/10 6 years.
720. Erez, J., K. Takahashi and S. Honjo. In-situ dissolution
experiment of Radiolaria in the central North Pacific Ocean.
Earth & Planet. Sci. Lett. V59, p. 245-254.
Demonstration that temperature is the main factor
controlling the rate of opal dissolution in the sea.
The results suggest that opal must dissolve at a rate
100 times faster in warm surface water (25°C) than incold deep water (1.5°C).
721. Ledwell, J.R., 1982. Gas exchange across the air-water interface. Ph.D. thesis, Harvard University, Cambridge, Mass.
Demonstration through wind tunnel experiments that thegas eschange depends on the square root of the molecular diffusivity.
722. Measures, C.I. and J. Edmond. 1982. Be in the water columnof the central east Pacific. Nature, V197, p. 51-53.
First measurements of the profile of Be in the sea.
723. Settle, D.M., C.C. Patterson, K.K. Turekian and J.K. Cochran.
1982. Lead preci~itation fluxes at tropical oceanic sites
determined from 2 °pb measurements. Jour. Geophys. Res. V87,
p. 1239-1245.
Estimates of the fluxes of 210 Pb and stable lead to theocean from measurements on rain water samples collected- at Einewetak, Samoa and the Florida Keys.
724. Carmack, E. and K. Aagaard. 1973. On the deep water of theGreenland Sea. Deep Sea Res. V20, p. 687-714.
Suggestion that subsurface of waters from two advective
regimes (via a double-diffusive process) leads to theformation of Greenland Sea deep water.
725. Swift, J.H., K. Aagaard and S-A Malmberg. 1980. The contribution of the Denmark Strait overflow to the deep North Atlantic. 'L'eep Sea Res. V27A, p. 29-42.
Demonstration that the water spilling through the
Denmark Straits forms at least in part at the surface
north of Iceland.
726. Lazier, J.R.N. 1973. The renewal of Labrador Sea Water.
Deep Sea Res. V20, p. 341-353.
Demonstration that deep water formation in the Labrador
Sea may occur only during winters of extreme severity.

677727. Ostlund H.G. and R.A. Fine.
1979.
Oceanic distribution andtransport of tritium.
In: Behavior of Tritium in the Environment.
International Atomic Energy Agency PublicationlAEA-SM-232/67, p. 303-314.
Paper discussing the distribution of tritium in the
Atlantic and Pacific Oceans.
72e. Enting, I.G. and G.l. Pearman. 1982. Description of a OneDimensional Global Carbon Cycle Model.
Division of Atmospheric Physics. Technical Paper #42, CSIRO, Melbourne,
Aust~alia.

In this model the ocean is b~oken up into two pa~ts(pola~ and equato~ial).
Deep wate~ is produced byt~anspo~ting su~face water from the polar unit to thebottom of the polar unit.
Pa~t of this water upwellsin the polar unit and part passes to the equatorial
unit where it upwells.
It includes the ~ain and destruction of carbon bearing particulates within thesea. The model is calibrated and tested using all theavailable natural and bomb-produced isotope distributions.
729. Hammer, P.M., J.M. Hayes, W.J. Jenkins and R.B. Gagosian. In
press. Exploratory analyses of trichlorofluoromethane (F-ll)
in North Atlantic water columns.
Profiles of freon-II in the Gulf of Maine, on the
Scotian shelf and in the Sargasso Sea.
730. Cline, J.D., R.H. Gammon and D. Wisegarver. In press. Meridional distribution of dissolved freon-II in the eastern North
Pacific Ocean.
_ ~omparison of the distributions of the freons withthose of salinity and tritium in the thermocline of thenortheastern Pacific Ocean.
731. Gammon, R.H., J.D. Cline, D. Wisegarver. 1982. Chlorofluormetanes in the northeast Pacific Ocean: measured vertical
distributions and application as transient tracers of upper
ocean mixing. Jour. Geophys. Res., (in press).
One-dimensional models for the vertical distributionsof freon-II and freon-12.
732. Broecker, W.S. and E.A. Olsen. 1961. Lamont radiocarbon measurements VIII. Radiocarbon, V3, p. 176-204.
Discussion of the ~ scale for the presentation ofradiocarbon results.
733. Urey, H.C:.
1947. The thermodynamic properties of isotopic
substances. J. Am. Chern. Soc.~ p. 562-581.
Early work of 13 C and 1 0 isotopes from the Cretaceous
belemnite.
734. Craig, H. 1957. Isotopic standards for carbon and correction
factors for mass-spect~ometric analysis of carbon dioxide.

678

Geochim. et Cosmochim. Acta., V12 p. 133-149.
Standardization of 13 C and IsO isotope ratios from theCretaceous belemnite. Secondary NBS reference CaC0 3
standard is also described.
735. Craig, H. 1961. Standard for reporting concentrations of
deuterium and oxygen 18 in natural waters. Science, V133,
p. 1833-1834.
Introduction of SMOW for reference standard of oxygenand hydrogen isotopes in seawater. Conversion of units
between PIE and SMOW is also given.
736. Friedman, I. and J. O'Neil. 1977. Compilation of stable
isotope fractionation factors of geochemical interest. U.S.
Geol. Survery, Professional Paper 440-KK, U.S. Government
Printing Office, Washington, D.C.
Conversion of stable isotope ratios units between PDB
standard and SMOW standard.
'737. Goldberg, E.D. 1974. The Sea, Vol. 5, Marine Chemistry.
Interscience publishers, Wiley and Sons, New York, 895 pp.
A series of articles on subjects in chemical
oceanography authored by experts.

679

FREQUENTLY USED CONSTANTS

Avagadro's Number (molecules/mole)
Ideal Gas Law Constant (cal/deg K)

6.02

x

10 23

1·99

Seconds per year
Minutes per year
Area of the Earth (cm 2 )Mass of the Atmosphere (grams)

5.1 x 10 21

Area of the Oceans (cm 2 )Volume of the Oceans (liters)
Mean Depth of Oceans (meters)

3.6 x 10 18
1.37 x 10 21
3.8 x 10 3Continental Runoff Rate (liters/year)
Ocean Volume/Runoff Rate (years)

3.6 X 10 16
38,000Mean Oceanic C02 Piston Velocity (m/day)
Mean Oceanic CO 2 Exchange Rate (moles/m 2 year)(at PC02 = 300 x 10- 6 atm)

2.8
18(for ~14C = 0 0/00)
Specific A~tivity (dpm/gC)
Halflife 14C (years)

1.18 X 10- 12
13.6
5730

(for T.D. = 1)
Specific Activity (dpm/liter H2 0)
Halflife 3H (years)

1.0 x 10- 18
7.1
12.4

5. 1

14 C/ C

3H/ H

680

x

10 18

DEFINITIONS OF ISOTOPE NOTATIONS*= [3He/4He)samPle -3 He /4 He )atm ] 100

3He/ 4 He) atm= [143Nd/144Nd)samPle _143 Nd /144 Nd )standard]

10,000143 Nd /144 Nd )standard
oD = [D/H)samPle - D/H)standard ]D/H)standard

=

1000

18 Q/16 Q)samP le - 18 0 /16 0 )standard]
18
16)
1000
[
0/ Q s tanda rd
13 ;12 )C
C sample -=[13 /12 )]C
C standard 1000

13 C/12 C)standard
14 C/ C )sarnP le -[o14 C -

14C/C)standard ]1000

14 C/ C )standard

2 (o13 C + 25) ( 1 +o14 C )

1000

*For 13 C and 180 the standards are isotope ratios for the
carbon and oxygen from the Cretaceous belemnite used as a standard
by Harold Urey in his early work (733). As this material has been
depleted, a secondary National Bureau of Standards reference CaC03
standard is used in most laboratories (734).
It should be notedthat for measurements on sea water the standard used for oxygen(and for hydrogen) is SMOW (s tandard mean ocean water).
Becausethe CO 2 prepared for mass spec trome t ry from CaCO 3 does not havethe same oxygen isotope ratio as the C03= in the mineral (because of interaction with the oxygen in the phosphoric acid used
to release the C02) the PDB and SMOW scales are quite different(735). Care should be taken in converting from one set of values
to the other (736).
For helium the standard is helium from the atmosphere.
For
neodymium the standard is the 143Nd /144Nd ratios for average earthmaterial (337).
Finally, for radiocarbon, the standard is 0.95
times the 14C/ C ratio in the National Bureau of Standard's oxalic
acid reference standard (732).

681

ABBREVIATIONS
Accumulation rate
Advection rate
Alkalinity
Antarctic Bottom Water
Apparent Oxygen Utilization
Area
Atmosphere·
Calcite compensation depth
CentimeterCurie
De cay cons tant
Denmark Straits Overflow Water
Depth
Disintegration per minute
Eddy diffusi vi ty
Equivalent
Flux
Gibbs Fracture Zone Water
Gram
Kilogram
Kilometer
Labrador Sea Water
Liter
Meter
Microequivalent
Micromole
MicronMililiter (CUbic centimeter)
IV1ili vol t
Minutel\1oleMolecular diffusivity
North Atlantic Deep Water
Percent
Per milPotential density
Potential temperature
Rain rate of particulate debris
Residence time
Revelle factor
Salinity
SecondSolution rate
Standard Temperature and Pressure
Surface water
Thickness of stagnant film
Tri tiurn Unit
Two-Degree Discontinuity
Upwelling velocity
VolumeWeddell Sea Bottom Water

682

A

w
Alk
AABW
AOU
A

atm
CCD
cm
Ci
A

DSW
z
dpm
D

eq, equiv
F

GFZW
g, gm
kg
km
LSW
1
m

}.leq
}.lmol, }.lm}.lml (cc)mV
min
mOl, m
K

NADW
%
0/00

a

e
R
T

R

S

s, secS

STP
swz
TU, tu
TDD
w, W
V

WSBW

Advection, 317
Alkalinity, definition, 67
Anions, control mechanism, 287
Antarctic bottom water (AABW), 337
Anthropogenic tracers, 383
AOD:

definition, 131
distribution, 131
Aragonite, solubilities, 63-64
Argon-39, deep sea ventilation, 377
Authigenic mineral, 45
formation of, 295
Barium:geographic pattern, 30
mean lifetime, 21
removal, 11
Benthic boundary layer, ~8Biogenic sediments:
calcite, 47
opal, 47Biointermediate constituents, 6,8,22
Biolimiting constituents, 6
Bioturbation, sediment, 257,562
Biounlimited constituents, 6,8Boundary stagnant film:
sea-air interface, model, 113
sediment-water interface, 86-88
Bubble entrainment, 127,130
Buoyancy flux, 329
Cadmium, distribution in the sea, 216,218,222
Calcite:biogenic, 47
changes, glacial-interglacial, 456
compensation depth, 100
degree of saturation, 61
distribution in marine sediment, 58
pore water resaturation time, 84
solubilities, 63
solution, bioturbation model, 487

683

solution, controlling factors, 84
solution rates, 269
storage, compensation time, 456-457
Carbon:control mechanism, 280
fixation, ratio organic to CaCD3, 476
reservoir sizes, 455
Carbon-13:
fractionation, 306
glacial-interglacial, 451-454,473
marine geochemistry of, 306
phosphate changes, 466
planktonic forams, 470
tree ring, 538
units, 306
Carbon-14 (see Radiocarbon)
Carbon dioxide:
anomalies, equatorial Pacific, 158
anthropogenic production, 500,501
atmospheric time history, 535
concentration in surface water, 149-150
equilibration times, 155
excess, equatorial surface water, 430
fossil fuel (see fossil fuel C02)
future levels, prediction, 550
ice cores, 445
potential partial pressure, 161,443
release, forest & soil, 502-503
solubilities in sea water, 111,112,151,152
Carbonate chemistry, 69
Carbonate ion, variation, 62
Carbonate saturation:
horizon, 71
spacial variation, 71
Carbonic acid:
1st apparent dissociation constants, 151,152
2nd apparent dissociation constants, 151,152
Calcium:
distribution, 11mean lifetime, 21
Cations, control mechanism, 289
Characteristic penetration depth, 319
Chemical composition:hydrothermal vents, 292,294
river and sea, 26-27, 290
saline lake, 293
Circulation pattern, deep water, 32,33
Conservative property, 334
Continental runoff, rate of, 247
Copper, distribution in the sea, 216,218,222
Dating:beryllium, 264
magnetic reversals, 267

684

potassium-argon, 264
radiocarbon, 255
uranium-helium, 312
uranium series, 257
Deep water, sources, 334
Denmark Straits Overflow Water (DSW), 335
Density gradient, 327-331
Detrital material, 45
Detritus accumulation rate, 45
Diapycnal mixing, 322,324
Diffusion, deep water, 317,319
Distribution coefficient, 230
e-folding time, 457
Element budget, earth's crust, 166,167
Element concentrations, river and ocean, 26-27, 290
Element, hydrolysis constants, 231
Element, removal times, 231
Film replacement model, 118
Flushing time, deep sea, 370
Foram, solution index, 98
Fossil fuel, reserves, 506
Fossil fuel CD 2 :atmospheric levels, prediction, 550
calcite solution, 552
oceanic uptake, capacity, 511
oceanic uptake, kinetics, 524
oceanic uptake, numerical model, 526,529
production, 500, 501
Gas exchange rate, 113
Gas flux,-116
Gases:
diffusivities in sea water, 119
equilibrium concentration in surface water, 114
solubilities in sea water, 112, 113
volume partition coefficient, 112, 113
Gibbs Fracture Zone Water (GFZW), 337
Helium:excess, 147hydrothermal vents, 295
Helium-3:
anomalies, 149
distribution, deep Pacific, 331
distribution, thermocline, 432
Hydrothermal vents, 292,294
Igneous rocks, chemical composition, 167
Inorganic carbon, total dissolved, 67
Input rates, river to sea, 25
Iron, transport in the sea, 226685

Isopycnal mixing, 322,324
Isopycnal surface, 322,391,392
Isotope fractionation, carbon, 236
Isotope ratios in reactive metals, 221
Keeling fraction, 514
Labrador Sea Water (LSW), 337
Lead:distribution in the sea, 218-222
isotope ratios, 223,226,227
Lead-210:depth profile, 174
distribution, 189
excess, sediment, 562-563
fallout rates, 191
production rates, 191
scavenging time, 189,199
Lysocline:
definition, 75
glacial-interglacial, 485transition zone, 89Manganese, transport in the sea, 226
Manganese nodules, 98
Metal transport, model, 205
Mixing rates, deep sea, 325
Model:
advection-diffusion, sea water, 318
biolimiting constituents, 15
stagnant film, 113
Mole, definition, 17
Molecular ~iffusion coefficients, 116,119
Neodymium, isotope ratios, 223-226
Nickel, distribution in the sea, 216,218,222
Nitrate:control mechanism, 284
consumption (denitrification), 142
distribution of, surface water, 47
geographic pattern, 31
Nitrous oxide cycle, 139
Northern Component Water (NCW):
definition, 335
conservative properties of, 340, 341
feed for, 345
radiocarbon in, 343
Nutrient gradients in the deep sea, 312
Ocean chemistry, glacial-interglacial, 467
Oeschger model, 526
Opal:biogenic, 47
diffusion coefficient, pore water, 55

686

distribution of, 47
rain rate, 56
resaturation time, pore water, 55
solution rate, 49,56
Organic carbon in sediments, 47
Organic matter, chemical composition of, 8
Oxygen:concentration, surface water, 127
consumption, respiration, 134
control mechanism, 284
deficiencies, deep water, 131
deficiencies, surface water, 131
excess, photosynthesis, 127geographic pattern, 31
glacial deep water, 484
Oxygen-18:
in forams, 446-447
Outcropping isopycnals, 322
Paleolysocline, 98
Paleoocean:chemistry, recorder of, 301
indicator, 301,311
Particulate matter, composition of, 13
Partition coefficient, 384
Phosphate:changes, causes of, 479
changes, glacial-interglacial, 466
control mechanism, 275
Phosphorus:
input rate, 25mean lifetime, 20,21
residence time in mixed layer, 20
scavenging time, 175
Piston velocity, 121Plant productivity, 130
Plutonium, distribution, 203
Polonium-210, distribution, 198
Potential temperature, 3,111
Potential C02 partial pressure, 161,443
Preformed values, 131
Protactinium-2jl:
231pa/230Jh ratio, 184,189
rain rate; 186,187
Radiocarbon:
14C/C ratio definition, 245
cycle, deep sea, 368
decline, fossil fuel, 544
distribution in the sea, 245
production, temporal variation, 374
production rates, 236-238
production, sun-spot, 544
Radiocarbon, bomb-produced:
distribution, thermocline, 412

687

Termination I, II, 454
Thermocline:
definition of, 15hydrology of, 388
Thorium:
activities, 170
isotopes in the sea, 172-184
scavenging times, 175
228 Th /L28 Ra ratio, 175-179
Thorium-230:
distribution coefficient, deep Pacific, 181
profile in sea water, 180
rain rate, 186,187
Trace metal, distribution, 215
Tracers, diapycnal and isopycnal mixing, 324
Tracers, purposeful, 438
Tritium:delivery mechanism, 384, 385
delivery rates, 386
distribution, thermocline, 399
penetration depth of, 402,403
temporal trends, 402
vertical distribution, 409
Two-box model, 15,240
Two-degree discontinuity (TDD), 359
Upwelling:
distributiton of, 158,319
bomb 14 C, 422
carbon dioxide anomaly, 427
evidence for, 397
rates of, 425,429
Uranium content:of coral, 311
of saline lakes, 311
Uranium series, activities, 170
Ventilation of deep sea:source areas, 312,317
deep Atlantic, 359
deep Indian, 365
deep Pacific, 365
Vertical mixing, rate of, 236-243
Vital effect, 468
Volatile elements, distribution of, 288
Weddell Sea Bottom Water (WSBW), 353
Zinc, distribution in the sea, 216,218

689

distribution, surface water, 415
time history, air, 389
time history, surface sea, 389
Radium:marine geochemistry, 13
226 Ra , distribution, 200-203
226 Ra , flux from sediments, 204
226 Ra / Ba ratio, model distribution, 245
228 Ra , deep sea, 326
Radon-222, deep sea, 326
Red clay:definition, 45
chemical composition, 167
Redfield ratio, 9Removal model, radionuclides, 171
Residence time, definition, 25
Revelle factor, 158,519
Salinity:
controlling factors, 289
changes, glacial-interglacial, 448
Saturometer, 80,83
Scavenging constant, 171
Scavenging times, 172
Sea salt composition, 2
Sediment mixing, 562
Sediment recycling, model, 212
Sedimentation rates, methods, 253
Selenium, distribution of, 22
Shale, chemical composition, 167
Silica: (see also Opal)
atom ratio, mean sea water, 11
control mechanism, 280
distributlon, surface water, 47
geographic pattern, 30
N-S vertical sections, 36-38
Solubilities of gases, 111,112
Southern component water (SeW), 349,359
Stagnant film:model, sea-air interface, 113
sediment-water interface, 562
thickness, radiocarbon method, 118
thicknesa, radon method, 122
Steady state, definition, 17
Strangelove ocean, 466, 472
Strontium, 13
Strontium-90:
corals, 412,413
delivery mechanism, 385
deposition rate, 386
Suess effect, 544
Sverdrups, 382
Temperatures, glacial-interglacial, 446

688

FOLOOUT CAPTION
The .foldout map shows the tracks followed during the Geochemical Ocean Sections Study (solid lines) and the NORPAX Hawal1Tahiti Shuttle (dashed lines). These expeditions provided much of
the data discussed in this book.
The GEOSECS Atlantic Ocean expedition was conducted during the period July 1972 to May 1973;

the Pacific Ocean expedition during the period August 1973 to June
1974 and the Indian Ocean expedition during the period December
1977 to March 1978.
The points on the GEOSECS tracks are stations.
Key station numbers are also given (the Atlantic stations
go from 1 to 121, the Pacific stations from 201 to 347 J and theIndian Ocean stations from 414 to 454)1
The Hawaii-Tahiti leg on
which the PeO ' te02, alkalinity and
4C/ C data discussed 1n thez
text were obtained, took place 1n April 1979.
The locations offi ve key deep water passages discussed in chapter 7 (1. e., theDrake, Vema, Romanche, Samoan and Atlantis II) are shown.
The
arrows making the locations of these passages designate the flow
'direction of bottom water.

690

,r
!1.
I.~