N 66 RP— |
NAVAL OCEANOGRAPHIC OFFICE REFERENCE PUBLICATION 1

BASIC ACOUSTIC
OCEANOGRAPHY

APRIL 1975

APPROVED FOR PUBLIC RELEASE;
DISTRIBUTION UNLIMITED

DEPARTMENT OF THE NAVY
WASHINGTON, D.C. 20373

PLIPVY OS
Al RI

MUNI

oo045175 3

iil

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MONI

0 0301

BASIC ACOUSTIC
OCEANOGRAPHY

MORRIS SCHULKIN

APRIL 1975

APPROVED FOR PUBLIC RELEASE;
DISTRIBUTION UNLIMITED

NAVAL OCEANOGRAPHIC OFFICE
WASHINGTON, D.C. 20373

NOO RP-1

a

FOREWORD

This publication has resulted from a 2 1/2-hour lecture with the
same title that the author has presented as part of the Catholic University
of America summer graduate course in Ocean Acoustics under the directorship
of Dr. Frank A. Andrews. Its purpose is to introduce the vast subject of
oceanography in an elementary and compact way, emphasizing those features
that would be useful and interesting for acoustic applications. The
term "acoustic oceanography" as used here includes studies of the ocean
using acoustics as a tool. However, the extremely limited scope of the
lecture has led to the inclusion of only one or two acoustic examples
under each chapter heading.

Jey hie A S
Captain, USN
Commander

iii

ACKNOWLEDGMENTS

The author is pleased to acknowledge the permissions granted by
Charles E. Merrill Publishing Company to use figures 1.1, 2.5, 2.6,
6.4, 7.4, 8.1, and 8.2 from "Oceanography" by G. M. Gross; by John
Wiley and Sons (Interscience Publishers Division) to use table l, p.
313 of vol. 1 and fig. 1, page 47 and fig. 1, page 657 of vol. 3 from
"The Sea" edited by M. N. Hill and fig. 3-4 from "The Earth and Its
Oceans" by A. C. Duxbury; by Prentice-Hall to use table 33, page 166,
fig. 183, page 669, and fig. 209A, page 740 from "The Oceans" by H. U.
Sverdrup, M. W. Johnson, and R. H. Fleming; by Houghton-Mifflin
Company to use fig. 7-13 from "Investigating the Earth;" and by the
University of Chicago, Department of Geography to use two of the above
figures plotted on its base map. The author also wishes to thank Ms.
Pat Jean of NAVOCEANO for carrying out the task of technical editing
and seeing the report through the arduous day-by-day problems of
getting it published.

iv

CONTENTS

loreratienetsl GHASCOOO DODO OCODOUDDOUO OOD OU OUOUDODUOODUOOUUG ataretelohstetenelions
Il6 Wbahertoyaoyorealoyy oho cond OO OGUOODOODOOUOOOUDGUGOE eeehcvonetelohelie el venetele: =
Deen SUbMawrine GOO Py Mejeiercisliccseie ls cretslie eel! el enelle) elchelicwelcleleiiel slic co « sic s 50
a. The Earth Beneath the Oceans ......ccccccccccsccce 360 60
Dem BathyMe tery mereverereseicierchenel sieteusicievelave: sherelelaveccielevelelerchelc sisters! stone ok
ComeBOEEOMES Cada Mets Sierecics clerelsie ce oleiie! cllel cl/eifeley oie) alisllelrelcjelctisleliclelclione oie
d. Bottom Acoustic LOSS ......-.cccecece shel cher such sLerererel shetiorcl sie) s
Bio wkeheatinyay Ihtonlfoyay GogaogdoGouU GOO UOOOU alekerencvalsielexetelclenoicichere siohevonenens
a. The Sea as a Biological Environment ........-.ccecee~ at
b. Organisms and the Composition of Seawater .......... O09
@o Wenesleya ISLECY 6 GG 6c odo OOOO CU OOO OUUUDD OD OU DDUOODODOUGOOUOG
d. The Deep Sound Scattering Layer ........... SOO O50 HO GROE

4. Chemical and Physical Properties of Seawater ........eceeeee-

ao Composmtdion) sn. si cciere siecle vies! « elelepelicvsta'c|elelehelolsnere eleveheiete ehelersicliele
b. Equation of State and Thermal Properties .........-.00.%
c. Optical and Radio Properties .......... HOO CODECS 5.05000
ime SOUS PC Oder ie.crsrstcieteiencherctenche oie lcheletielaiere elciicrsclc\ sleyeyensteyevenever eile
elon HSOUNAWADSOF PEON eye <eye cl scl at o1cic whe ce wee alot sieheferehoterels SodG050C

DewwlheSteadyuStatero& (the Ocean acc cess ce elec s cie's cre cle slelstejels sc)

a. Heat Transport by Winds and Currents ...... elctcheletevets S000
b. Three-Layered Vertical Temperature Structure ..........
c. Salinity and Water Masses ....... slicliel stolelistevelete shalianotavonovelevers

d. Sound Speed vs. Depth Structure ........c.ccsccsccscess
e. Modes of Sound Propagation in the Ocean ......ecccccees

Ge Waviesmin the Occambacttteisieccresictoiets, clstererctoronerchototonele olecarevene Sictevetercre

a.“ The Sea ‘Surfacetand Acoustic HELECES inci cose sc clec «0 wc

b. Energy Spectrum of the Wind Waves ..c.eccesccessecceace
c. The Thermocline and Internal Waves ...cesscccessecs Soo
d. Acoustic Propagation and Internal Waves ..... adotebncad
Biblio geraphy, ceisieie cle o'eis cecil’ < cle) sl eveleicleieleleiete Polen stevorevelel steric atateieteheneleteilere

FIGURES

Page

Submarine Geology
Layers of the earth and their densities ........ eiehelafetencney-tetencren: 23
Earth's crust consists of units moving as rigid blocks ....... 24
Formation of oceanic crust at ridges and destruction

UT GEGETICHES @ ctolersieiels isle cfeve sl evel eielisic sc iaie sls atepetel aislsloxolaiefelelcl ole sheer 2D)
Bottom profile of the North Atlantic Ocean ..............0e0- = 26
Bottom topography chart of the North Atlantic Ocean .......... 27
oxy Inbigoysyoywera lope (Gta Mano FOO mOOOo OOOO OOOO OO OOGE DMO UGS 28
Distribution of deep-sea sediments ......... Brel evayaie ouebe: srevensiaye evens 29
Typical method of continuous seismic profiling ............06. 30
Bottom loss for “Hast! and) USilow.) Sediments) yie.se) oe ciels sels evels elon 31
Marine Biology
The classification of the benthic and pelagic marine

ETUValst OME Mtammoneh apopenetetevetev honey slater chcVenevensiebey Lorene ctenel er HOGUUHOOGUOOG 32
Transparency chart of the North Atlantic Ocean ............ oc Ss.
A plankton sample with representatives of the meroplankton ... 34
Representation of the "deep scattering layer" ........eeeeee. 7 35
Variation of effective scattering area with wavelength for

small solid particles and bubbles ...... Melotorevorcnebetelercicuckerstenete 36
Representative fishes possessing swimbladders ............. co (oY

Chemical and Physical Properties of Seawater

Dissolved chemical constituents of Seawater ...cceccccececvcee 38

Detiniteron, of Sadeimd ty) vrei eiere w sie rene Ri ay aional etek vol ohce cuoheucnerszcvcber ey cusvoneters 39
Gases in deep ocean water as compared with their

CONCENEEAELONG Tm walie ayeveveyelereisie clere cycle) ere cloteie)erele ere ayeietclerstcloteNeraye 40
Absorption of waves in seawater ..... sre) (site leilevieyeleleveieiele el exes lets e soo (lh
Speed of sound in seawater (Wilson) .....ccccesecsvccccnnccces 42
Magnesium sulphate complex in solution ..... Dos oDGGOON OOD OOO 43
Absorption of sound in seawater (Schulicinvarsn), OO CO OOOO Ie 44
Absonptionvof Sound: inyseawater lUhOGp)) -iewcicte co eicletereie sicrels ale 6 45

The Steady State of the Ocean

Radiation balance in the northern hemisphere and average

ocean surface temperatures ......ceeeeeee SaudHbUgOoOORDOOAND 46
A schematic representation of the cellular circulation

of the atmosphere on a rotating earth .... cece ccccccccces 47
Water movements in a wind-generated current in the northern

hemisphere ....... aionoielevertatelictcvehe eh eustenetererarstenstetete ones atop etieken eke we AS
Generalized atmospheric and oceanic circulation .......... Good) «Ue,
Schematic diagram of ocean layered structure .........eeeeece - 30
Representative water temperature profiles ......cccsecccsescce Dil
Wind and convective mixing ...... So poaddndododenUGsoUoOuOUGD coo Sy

vi

FIGURES (con. )
Page

Percent occurrence of layer depth in North Atlantic

DAES CAS OMe raleoteleteMe Nero ebencoier ticle elo olf Melatel hele eclicrelioteltejie/si <rlelle\isyie:.n1.<1'=\\s (e/eil= 53}
The change in average midocean surface salinities in

response to latitudinal. changes in evaporation and

jpxaealaliehenllony Godumcoonddo 06000 DODO udGOUUO oO OOUu OOOO oUO oo co 54
Temperature and salinity of the water masses of the

Nellemen@ Oey oodobonoabboonooQbooodbshooodogoUU Oo OUOUb GUO 5)5)
The geographic distribution of water masses that have common

vertical salinity and temperature distributions ..........-. 56
Basic temperature and sound velocity structure of the

QCD. Gasicsobdd Gob osn so cameo obEnEOboo OOD A Doo GoUO a KOO ooo OUD 3)7/
Sound channel axis depth vs latitude (55°W longitude) .......- 58
Sound) transmission modes!) (schematic)! 7...).:.. 2 ss 0rs oe cle) elerele wiels s) «ele 3)8)

Waves in the Ocean

Saimpillemsecamwavermandwest Srp t Sie cierelic sielere cre sc elelcreleielsiere sjenelisiejeielel s« 60
Wave profile and water particle motions ......c..eecceeecceees 61
Streamlines of flow in finite amplitude capillary waves ...... 62
Speed of gravity waves and capillary waves as a function

Of mth eaicswaviel: Smet higtinerare ts tojci oo levees s/atlsifeietrorielleroters) el eNedeter's ereteneis-(o es. ¢ 63
Seal sitate code: \emc. e «seters ais FROG COCO OOD OODGO ODEO LOD OOS UDO bOe 64
Deepwater ambient jmoitse Mevieli’s i cic so dice eles creel e sic ere cllete ateiarshenepe 65
Sea state and acoustic loss at the sea surface ..............- 66
Comparison of an observed wave spectrum to empirical

SPeCiral Of PVE TS ON-MOSKOWALE Zits tencletelcteleleieiele laiie stelleieieleevekenonelishetonere 67
EsSotherms in anternals wave! MOEWOMS) fie <iecc olersie sie) oils fore10) eheieselekeye) <= = 68
MMS rama le WAVIe) ISPCCEIGA) Were chereiiel<hels)/e//<1 01 «1:e)(e1.« folie tole! ele! sje cl eda)levie s'olelor sie iele\ overs 69
Diagram of rays in the sound medium, which has an internal

WAVIEMOMMEMe ithe GMO! Maliryes a clletey sl eles; sles «ici sis) at elves eve! oeseieloiolslenseyelessieve 70

vii

1. Introduction

Topics chosen for presentation in a single lecture on the study of
the ocean using acoustic signals must be carefully selected and
treated briefly. The ocean is mostly opaque to light and radio waves,
so that acoustic and other mechanical probes become extremely
important in gaining information about the insides of the ocean. The
generalized view taken here of acoustic oceanography is one of many
possible approaches.

The oceans cover almost 71 percent of the Earth's surface and have
a mean depth of 3,800 meters. Some ocean trenches are deeper than the
height of the highest mountain on earth. In this report submarine
geology is treated first. Natural processes occurring beneath the
ocean floor influence the shape and structure of the floor. Sedimen-
tary processes also modify it. The resulting morphology affects
acoustic propagation when the ocean floor is involved. Next, marine
life is studied in terms of interaction with the ocean environment.
Use of acoustic techniques has revealed some interesting habits of
sea life, of which one example is the diurnal migration of the deep
sound scattering layers. The section on chemical and physical
properties of seawater is slanted towards speed of sound and absorption
of sound. Although the ocean is a variable medium both temporally and
spatially, the next section describes mean state with regard to
temperature, salinity, and currents. Acoustic effects caused by
stratification of the ocean, as represented by the surface mixed layer,
thermocline, and nearly-isothermal cold deep ocean, are treated in
terms of the surface duct and deep sound channel. Examples are given
of the four chief paths of acoustic wave propagation using ray sche-
matics. The final topic concerns ocean surface waves and internal
waves and their statistical variability. Acoustic effects related to
the ocean surface, such as noise spectra and sea surface loss, are also
treated briefly.

2. Submarine Geology

Submarine geology concerns topography of the sea floor, composition
and physical character of sedimentary and igneous materials found on
the ocean bottom, and processes involved in the development and modifi-
cation of topographic relief.

a. The Earth Beneath the Oceans

Earth is a sphere of 6,371-km radius and is considered to be
composed of four idealized concentric spherical layers: crust (0-33 km),
mantle (33-2,900 km), liquid outer core (2,900-5,100 km), and inner core
(5,100-6,371 km) (fig. 2-1).

The boundary between Earth's crust and the denser mantle at an
average depth of 33 km is called the Mohorovicic discontinuity, or the

Moho for short. The Moho lies at a greater depth under land masses

1

(20-50 km) than under oceans (5-10 km). Earth's crust consists of
several large rigid plates or blocks that float on the more plastic
mantle. These plates move relative to one another and most, if not
all, earthquake and volcanic activity occurs along plate boundaries.
Mid-ocean ridges represent boundaries of plates that are spreading
apart. Peripheral mountain ranges, volcanic arcs, and deep-sea
trenches represent the boundaries of adjacent plates that are con-
verging (fig. 2-2). When the strength of the Earth's crust is
occasionally exceeded locally, sudden ruptures occur that displace
rock masses along fractures called faults. Most faulting occurs at
the borders of crustal plates.

Many Earth scientists believe the lateral spreading apart of
crustal plates is caused by motion of convective cells in the mantle
(fig. 2-3). Molten material in these cells rotates about cell axes,
forcing hot basaltic mantle material up through and along the underside
of the crust, slowly dragging the crustal plate until a zone of down-
ward cell motion is reached. The Mid-Atlantic Ridge is estimated to
be spreading on the average of 1.25 cm/yr, whereas the East Pacific
Rise is spreading at a rate four times as great. Since new crust is
forming, old crust must be destroyed somewhere because the earth is
not growing larger at the required rate. A process of subduction occurs
at trenches, which are depressions where crustal plates are forced,
or sink, into the underlying mantle. Slippage of material into trenches
gives rise to earthquakes and volcanic activity, which leads to forma-
tion of seamounts and island arcs.

b. Bathymetry

Most detailed measurements of ocean depths and submarine
topography have been made in the last 50 years, since the development
of echo-sounding equipment. This acoustic equipment measures travel
time of a sound pulse to the sea bottom and return, producing a
continuous trace of the profile of the sea floor along a ship's track.
Figure 2-4 shows a bottom profile of a vertical section of the North
Atlantic Ocean along a track between Cape Henry, Virginia, and Rio de
Oro, Africa. The continental margin of each continent consists of a
shelf, a slope, and a rise. Ocean basins are separated by the Mid-
Atlantic Ridge. The old volcanic island of Bermuda is also seen.

After collecting and assembling appropriate bathymetric data
from many such tracks, detailed bottom contour charts are prepared
which may be used for navigation as well as for other scientific pur-
poses. Figure 2-5 is a bottom topography chart of the North Atlantic
Ocean. This chart shows the lateral extent of the major physiographic
provinces of the ocean floor.

The hypsographic curve (fig. 2-6) is a plot of the cumulative
areal distribution of heights above and depths below mean sea level.
This curve is used to show the area of different sea floor zones of
the Earth's oceans. A shallow zone at depths from 0 to 200 m occupies

2

about 7.6 percent of the total ocean area. In this zone are the conti-
nental shelves, the shallow aprons surrounding continental land masses.
Continental shelves vary greatly in width, from a few miles off the
California coast to some 800 miles in the Siberian Arctic. Although
the average slope of the shelf is gentle, within the shelf boundaries,
terraces, ridges, hills, depressions, and steep-walled canyons may
occur.

Next, a deeper zone of the continental margin from about 200
to 3,000 m deep takes up about 15.3 percent of the oceanic area.
Within this zone are the continental slopes of the oceanic floor,
regions in which continental land blocks drop off rapidly to the
deeper sea floor. A common feature of the slope is the occurrence of
extraordinarily large gullies, known as submarine canyons, which are
steep-walled fissures penetrating the slope and cutting across the
shelf.

The continental slopes, with a typical gradient of about 4°,
drop smoothly to the incline of the continental rise which has a
gradient of about 0.3° to 0.05°. The continental rise then merges into
the floor of the ocean basins.

The greatest portion, 75.9 percent of the sea floor, lies at
depths ranging from 3,000 to 6,000 m, the zone of large oceanic
depressions called basins. These are usually considered to be the
Atlantic, Pacific, and Indian. Other large areas also are sometimes
considered as separate units, such as the North Polar Basin, North
American Basin, and Pacific-Antarctic Basin.

A major feature of the ocean floor is the mid-ocean ridge
about 1,500 to 2,000 km wide that extends for 60,000 km and covers
about 25 percent of the ocean floor, an area about equal to that of
all the continents. Throughout most of its length, the ridge is split
by a deep rift in which many earthquakes originate, the ridge being
the locus of a crack in the Earth's crust. The mid-ocean ridge runs
nearly twice around the Earth, occupying about the same position on the
hypsographic curve as continental slopes.

Only 1.2 percent of sea floor lies at depths ranging from 6,000
to 11,000 m. This is the area of troughs, trenches, and deep basins.
The deepest trenches are found in the Pacific Ocean and exceed 10,000 mn,
whereas the deepest point in the Atlantic Ocean is the Puerto Rico
Trench at 8,750 m. Note that the depth of the deepest trench is greater
than the height of the highest mountain.

c. Bottom Sediments

The crust of the ocean floor is coated by a sedimentary
blanket that has buried or modified the original igneous surface. The
morphology of the sea floor is, therefore, the result of generation of
igneous topography through processes of faulting and volcanism at the
crest of the mid-oceanic ridge, lateral motion away from the ridge
crest in a conveyor belt fashion, and slow burial by sedimentary
processes.

3

Regional elevation depends on age of the particular parcel of
crust. The original igneous crust is mantled by sediments into zones
reflecting age of the basin, proximity to continents and sediment
sources, and areas of biologic productivity. The three primary
sedimentary processes for smoothing of the original igneous rock are:
(1) continuous pelagic "snowfall"; (2) gravitational slumps and slides,
including turbidity flows from both terrigenous sources and, on a
smaller scale, the elevated undersea peaks; and (3) redistribution of
sediment by bottom currents.

Figure 2-7 shows the geographic distribution of deep sea
sediments by source as explained below. Sediments may be classified
in increasing size order as clay, silt, sand, and gravel. Sediments
are also classified by source.

(1) Terrigenous -- Organic or inorganic land-derived materials;
material derived from inorganic rock is lithogenic.

(2) Biogenic -- Organic or inorganic material from marine
organisms.

(3) Volcanic -- Submarine or airborne volcanic material (lava,
ash).

(4) Hydrogenic -- Materials crystallized from seawater, such

as manganese nodules.

(5) Cosmic -- Materials of extraterrestrial origin such as
cosmic dust and tektites.

Sediments are often classified by the biotic form that produces
their major biogenic component. These sediments are called oozes.
Microscopic free-floating plants and animals of the sea leave hard
remains or tests of either a calcareous or siliceous nature. Calcareous
deposits, though widespread, are restricted to shallower regions of
the pelagic zone than siliceous deposits, since calcareous remains
are more soluble than siliceous remains. Brown or red clays predominate
in deeper parts of the ocean basins under surface waters of low
biological production. These clays are derived from the very smallest
particles of wind-carried atmospheric dust from land.

Coring of the sea bottom and collection of material by use of
dredges and grab samplers are techniques for obtaining samples from the
upper layers and surface of the sediment. These methods are slow,
laborious, and expensive. Deep sea photography is also used for studying
the characteristics of the sea bottom. However, seismic profiling in
the deep ocean, introduced for routine shipboard use in 1956, has
become one of the most widely used tools of marine geologists and
geophysicists. Seismic profiling is a method of studying sediment
thickness and layering, using the penetrability of low-frequency acoustic
energy into the bottom (fig. 2-8). If the acoustic energy density is

Th

high enough, the energy remaining after partial reflection at the

sea floor penetrates sediments and suffers partial reflections at
each successive boundary. The bedrock under the sediments also
reflects energy, marking the bottom of sedimentary material. Sound
travels at a determinable speed in water and in sediments; the time
interval between the outgoing pulse and return of the reflected
energy then determines distance between reflecting horizons. Subbottom
profiling yields information on the sediments, topography, and
structure of the oceanic crust. For determining crustal thickness
with depths of many kilometers, sound in the frequency range of 3 to
30 Hz is used; whereas for measurements of sediment thickness of the
order of tenths to a few kilometers of sediments, seismic reflection
profilers in the 30- to 500-Hz range are used. For studying the
detailed stratification of the upper sediments, in the top 100 n,
echo-sounders in the range 3.5 kHz to 12 kHz are used. As can be
inferred from the previous information, ocean bottom acoustic charac-
teristics are strongly frequency or wavelength dependent.

From seismic studies it has been found that the average thick-
ness of the crust including the ocean sedimentary layer, down to the
mantle is 6.0 km, compared with an average crustal thickness of 33 km
on land. The average thickness of the present oceanic consolidated
or compacted sediments is 3 km. This is equivalent to 9 km of
unconsolidated sediment, which still has interstitial water in its
pore spaces. The thickness of sediments and rates of sedimentation
are used to estimate the age of an oceanic basin, with 200 to 600
years considered the average time for settling to a depth of 2,900 nm.
This corresponds to a particle size of about 4 x 10-5 cm, which is in
the silt-size range. Of course, this is a grossly oversimplified model.

d. Bottom Acoustic Loss

The loss of acoustic energy impinging on the bottom is quite
important for long-range sound propagation in the ocean, as well as
for understanding subbottom structure. Acoustic energy may be lost by
partial transmission into the bottom and subsequent absorption or by
scattering into nonspecular directions due to bottom roughness. Bottom
reflection loss can vary considerably from one location to another.

Important parameters are the sound speed and density contrasts
between water and sediment layers in the bottom as well as porosity of
the sediments. Thicknesses of the bottom layers must be considered
also, especially in terms of acoustic wavelength or frequency. For
sediments with porosities greater than about 65 percent, sound speed
in the bottom is slower than in the deep bottom water. For such
sediments there is an angle of intromission, at about 10° grazing
angle, with very large losses of about 26 dB. The losses then decrease
to about 14 dB from 40° to 90° grazing angles. On the other hand "fast"
bottoms, with sound speed greater than water, will show a critical
grazing angle at about 30°, with practically 0 dB loss from 0°-30°
grazing angles, and a sharp rise with losses of about 15 dB from 60° to
90° (Eig. 2-9).

3. Marine Biology

a. The Sea as a Biological Environment

The sea has several distinct advantages over the land for the
development of life. Water, necessary to all life, is immediately
available. Beyond this, seawater contains in solution the required
gases and other materials for growth, and is very similar in composition
and solute concentration to the body fluids of all animals. The result-
ing near equality of the osmotic pressure of seawater and body fluids
permits the direct intake of nutrients and excretion of waste in some
organisms. Buoyancy provided by the high density of seawater reduces
the need for skeletal structure, whereas the high specific heat of
seawater reduces the possibility of rapid temperature change. The
range of pressures experienced in the ocean as a whole is very great,
and many organisms are consequently restricted to specific layers, but
all life is not excluded from the deep abyssal regions. Bathyscaphe
dives, in particular, have confirmed the existence of marine life in the
deepest parts of the ocean.

Overall, a considerable variety of conditions exists in terms
of salinity, temperature, illumination, and pressure in the ocean.
Great variability with time is largely confined to coastal waters and
shallower layers. Here, marine forms must be adaptable to changes or
capable of adjusting themselves to the optimum condition. Light in
the surface layer is particularly important for photosynthesis by
floating microscopic plants. Notably, there is an adjustment to
lighting conditions by some marine life, resulting in a diurnal migra-
tion toward and from the sea surface. Large population of this marine
life can form sound scattering layers. As distinct from regions where
marked variability exists, large volumes of deep water have remarkably
constant conditions, and specific forms appropriate to the particular
environment can develop. Circulation processes in the sea are important
to life in many ways. These processes maintain oxygen content,
replenish the supply of nutrients, and aid in dispersal of waste material
and young life in the form of spores, eggs, larvae and young adults.
Clearly, however, unexpected and unseasonable disturbances of the
circulation pattern can, by moving life into an unacceptable environment,
result in considerable destruction.

Formally, the marine biological environment is divided into the
water region, called pelagic, and sea floor region, referred to as
benthic. The water region is subdivided horizontally into the neritic
zone, which lies over continental shelves (low tide to 200-m depth) and
the oceanic zone, which lies seaward of the shelves. The oceanic zone
is further divided into depth regions: epipelagic, from the sea surface
to about 200-m depth; mesopelagic, 200 m to 1,000 m; bathypelagic, 1,000
to 4,000 m; and abyssopelagic (hadal), the deepest parts of the ocean
(Ere S=1))%.

The most significant biologic zonation is the division into
the euphotic or photic zone, in which there is sufficient solar
radiation for active photosynthesis, and the deeper aphotic region.
The photic zone may extend to about 180 m in clear oceanic water at
low latitudes, or may be as shallow as a few meters in the more turbid
coastal regions of temperate latitudes (fig. 3-2).

There is considerable seasonal change in surface solar
radiation in the temperate zone; therefore, thickness of the photic
zone varies throughout the year. In the spring, when radiation levels
increase, plant life grows rapidly, using nutrients supplied to surface
watérs from depth by winter mixing. Surface heating in spring imposes
a stability on the water column that reduces vertical mixing and
increases the residence time of phytoplankton in the layer above the
developing thermocline. Retention of phytoplankton near the surface,
increased radiation, and available nutrients trigger the spring bloom.
In the tropics, low nutrient supply prevents high productivity of
organisms; in temperate regions, both light and nutrients may limit
production; and in polar regions, light alone controls productivity.
Nutrients are supplied to the photic zone by: (1) rivers which supply
land material, and (2) upwelling of water from depths below the photic
zone where nutrients have accumulated due to decomposition. In the
sea, as on land, photosynthetic plant life must exist in order for
animal life to survive.

b. Organisms and the Composition of Seawater

Seawater is more than a solution of salts in a certain proportion,
for marine life is an inherent part of the sea, and marine organisms
play an important role in determining the chemical and physical
properties at any particular location and time, e.g., they have an effect
on oxygen and carbon dioxide gaseous contents. Beyond this, certain
elements occurring in small concentrations in seawater are removed in
relatively large amounts by marine life for skeletal, structural, and
other uses. Inorganic compounds of nitrogen, phosphorus, iron, and
silicon -- the plant nutrients -- fall in this class, and their
distribution is affected appreciably by biological activity.

Life cycles of the various organisms in the sea are closely
interwoven. Plants of all sizes furnish the basic food for all life
and grow in the upper illuminated layer. These plants absorb carbon
dioxide from solution and release oxygen. They also absorb phosphates
and nitrogen salts and convert them into organic matter. However, the
inorganic substances that they use for growth must be made available
to them near the surface, in spite of downward movement of particulate
matter in the sea under the influence of gravity. Availability of
these materials at the surface results from circulation of ocean waters,
convective mixing, and upwelling from the bottom. The inflow of
minerals from land is a small factor in supplying nutrients for sea
plants. Whenever plant nutrients are abundant near the surface, large

7

marine populations occur, as compared with other parts of the ocean.
In deep water, such areas can be identified by their color, for water
rich in life is greenish, whereas "desert" water is blue.

c. Marine Life

The sea's inhabitants range from a multitude of microscopic
one-celled organisms to the 100-ton blue whale. Smaller forms,
especially plants, are extremely prolific under proper conditions and
furnish the food supply for successively larger forms of animals. In
describing the population of the sea, three major groups are
distinguished depending upon habitat: mnekton, benthos, and plankton.
Nekton is the name given to swimming animals that can move freely over
large distances. The term benthos refers to the sessile animals such
as sponges, creeping forms such as crabs, and burrowing forms such as
worms found at and in the sea bottom. The plankton includes all
floating plants (phytoplankton), animals that live permanently in a
floating state (zooplankton), and also larvae and eggs of animal
benthos and nekton (meroplankton). The plant plankton (e.g., diatoms,
dinoflagellates, and coccolithophores) and animal plankton
(e.g., globigerina and radiolaria) are so numerous that extensive
areas of the ocean floor are covered with oozes (fig. 2-7) containing
their skeletons. There is some question regarding the grouping of
motile dinoflagellates as plants even though they carry on photosynthesis.
Figure 3-3 shows a zooplankton sample, including meroplankton.

In the sea, as on land, plants are capable of developing com-
plex organic substances from the simple compounds dissolved in water.
Without these plants animal life would be restricted to a small region
near the shore where organic material of terrestrial origin exists.
Nearly all marine plants are thallophytes, primitive plants with no
true root, stem, or leaf. They include marine algae and fungi, and
bacteria are frequently included with them. Brown algae vary greatly
in size, and plants of this class form the "kelp beds" commonly
harvested for commercial purposes. Yellow-green algae are mainly
floating forms and include diatoms, which are unicelluar in structure,
and dinoflagellates, again typically unicelluar, but possessing two
flagella for locomotion. The latter are second only to diatoms as
producers in the marine plankton. Some of the latter have the capacity
to luminesce when disturbed. Mosses (bryophytes) and ferns (pteridophytes) ,
two major divisions of the plant kingdom, are not represented in the
sea. The highest plants, the spermatophytes, are represented, however,
by some 30 species of flowering plants. They have invaded the sea by way
of fresh water.

Of more than 50 animal classes, only four are exclusively
nonmarine. In contrast to plant life, all major groups (phyla) of
animal life are represented in the sea. Each phylum is subdivided
successively into class, order, family, genus, and species. When
thinking of animal life in the sea, one tends to emphasize the higher

order of marine animals such as the whale and larger fishes, but
these are an extremely small part of the total animal life.

d. The Deep Sound Scattering Layer

Volumes have been written on the complex sound production
and reception by porpoises and other cetaceans for food hunting,
navigation,and communication. Noise making by croakers and snapping
shrimp is also covered elsewhere. This section, however, will be
concluded with some remarks regarding the acoustic effects of the "deep
scattering layer.'' The deep scattering layer is named from its
property of scattering sonar signals and showing up as layers on depth
sounding records (fig. 3-4).

Actually, there appear to be two principal deep scattering
layers. The shallower layer has an average midday depth of 240 n,
with extremes between 185 and 405 m. Average thickness is about 75 nm.
The deeper of the two layers has an average midday depth of about 500 nm,
extreme values being 405 and 590 m. Average thickness of this layer
is 130 m. The average migration rate at sunrise and sunset of elements
in the 500-m layer is about 7.5 m/min and that in the 240-m layer is
about 4.5 m/min. The 500-m layer usually is not found in water depths
of less than 1800 m.

The food chain of the sea is intricately involved in biological
communities forming the scattering layers. A range of plant and animal
sizes occurs from microscopic to the largest fishes, all feeding on
each other. Scattering layers are sensitive to light, approaching the
surface at sunset and going deep at sunrise (fig. 3-4). Levels to
which the animals descend appear to be related to their negative
response to light and sensitivity to sudden temperature decreases.

Acoustic scattering properties of deep scattering layers are
similar to those expected from a collection of simple air bubbles in
water and show resonant peaks as a function of size and depth (fig. 3-5).
In fact, bubbles scatter sound much more strongly than solid particles
of the same size.

The main scatterers are fish with gas-filled swimbladders. To
maintain their constant size, these fishes are able to generate gas in
their swimbladders as they descend and vent gas through their gills
via the blood stream as they ascend. As verified by photographs and
net trawls, most numerous mesopelagic fauna,with swimbladders, between
depths of 150 and 1,000 m are myctophids or lantern fish, bristle-
mouths, and hatchetfish (fig. 3-6). Myctophids reach a maximum length
of from 4 to 12 cm, with swimbladder diameters of about 4 percent of
their length. Resonant frequencies are observed between 3 to 20 kHz,
corresponding to fish lengths of from 10 cm to 1 cm. It is interesting
that siphonophores have a small carbon monoxide bubble in their float
and scatter 12-58 kHz at 100 m depth and 24-111 kHz at 400 m depth.

9

Some indication of regions of relative sound-scattering
strength in the ocean may be obtained from figure 3-2, which shows
the geographic distribution of transparency, since transparency in the
open ocean would be affected by the density of plankton as well as
solid particles. The Sargasso Sea region, for example, is exceptionally
clear, whereas the upwelling regions near Iceland are rated only clear.

4. Chemical and Physical Properties of Seawater

a. Composition

Seawater is a dilute solution of salts. Its composition in
terms of ion concentrations is given in figure 4-1. Four ions (Nat,
Met, Cl”, and SO, ) constitute 97 percent of the total. The formal
definition of salinity follows: "Salinity is the total amount in grams
of solid material dissolved in 1 kg of seawater when all carbonate has
been converted to oxide, all the iodine and bromine have been replaced
by chlorine, and all organic matter has been completely oxidized"

(fig. 4-2). It is expressed in grams of salt per kilogram of seawater
(parts per thousand by weight of seawater is represented by the symbol
°/oo). An empirical relationship between salinity and chloride content
(chlorinity) is used to convert a measurement of chloride content to
salinity,

S °/oo = 1.80655 Cl ©/oo. (S = 35.00 °/oo for Cl = 19.37 2/o0)

Proportions of the different constituents are nearly constant
in the open sea, although the total salinity varies between about 33
and 38 °/oo, depending upon precipitation, evaporation, and ice melting.
Considerably lower salinities may be observed in coastal regions and in
almost-closed seas with low evaporation and a considerable inflow of
fresh water, e.g., the Baltic, where the salinity may be as low as 5
to 8 O/oo. Conversely, with little inflow and high evaporation, as in
the Red Sea and Persian Gulf, salinities reach as high as 40 °/oo.
Nevertheless, the assumption of a mean salinity of 35 °/oo for seawater
(90 percent of the ocean differs by less than 3 percent from this value)
is adequate for many applications where high precision is not required
and to define many of its physical properties.

Trace quantities of at least 50 of the elements have been
identified in seawater. Surprisingly, the sea is very poor in iron.
Significantly, copper replaces iron in the blood pigment of many marine
animals. Sea animals and plants sometimes concentrate trace elements
in their systems way out of proportion to the natural concentration.
For example, seaweeds can concentrate in their tissues the small amount
of iodine in the sea to the point where they may be harvested to obtain
commercial quantities. In some fish, concentration of heavy metal
pollutants in their systems makes them poisonous to eat. Considerable
attention has been given in recent years to the radioactivity of sea-
water, the radium content averaging about 0.07 x 10-12 0/00, which is
about 10-4 of that in rocks of the Earth's crust.

10

Gases may enter solution either by purely physical absorption
or chemical combination with the liquid. For the ocean-atmosphere
system, physical absorption occurs in the case of oxygen, nitrogen,
and rare gases. Carbon dioxide, on the other hand, reacts with water
and dissolved salts. In comparison with the oxygen-nitrogen ratio in
the atmosphere of 1:4, that of air dissolved in water is more nearly
1:2. Saturation quantities are a function of water temperature and
decrease by 50 percent from polar to equatorial regions. There is
ample air available at the sea surface, therefore, gas contents in the
upper layers are near saturation. As nitrogen is relatively inert,
only minor variations with depth are observed, since all water has
been in contact with the atmosphere at some time. The oxygen content
is markedly affected by biological factors, such as generation by
photosynthesis in the upper layers, which occasionally results in
supersaturation.

Oxygen consumption by animal respiration and bacterial oxidation
reduces the oxygen content and results in a pronounced minimum at
depths of 100 to 1,500 m in the middle and lower latitudes. Deeper
waters of polar origin are relatively rich in oxygen. In stagnant
water, where there is no photosynthesis and oxygen depletion occurs,
as in the deep portion of the Black Sea, hydrogen sulphide may be
formed. The partial pressure of CO2 in the sea in contact with the
atmosphere tends toward equilibrium with that in the air. However,
areas exist where carbon dioxide-rich water is brought to the surface
by rising currents, while in other regions the C02 partial pressure
is frequently low in spring and summer where plant plankton is
assimilating. Figure 4-3 is based on the condition that the volume of
the atmosphere is about 2.9 times that of the ocean. Figure 4-3,
comparing gas concentrations in air and seawater, shows that the ocean
is) ay sank tor C02).

The pH is a measure of hydrogen ion concentration on a logarithmic
scale. The pH of pure water is 7. An acidic solution has a pH of less
than 7, whereas an alkaline solution has a pH greater than 7. Seawater
is normally alkaline with a pH between 7.5 and 8.4.

b. Equation of State and Thermal Properties

The equation of state represents the density of seawater as a
function of salinity, temperature, and pressure. This equation is of
particular importance in oceanography in that minor changes in density
affect stability, water mass transfer, and turbulent mixing. For
these purposes, the oceanographer must know the density to 1 part in
10- and must resort to indirect measurements of chlorinity, electrical
conductivity, refractive index, or the like for density determination,
with corrections for depth and temperature. The expression for density
is unwieldy and in practice it is customary to use the notation

Oo = (p-1) x 103 where p is in g/cm3, and to use tables of Of as a
function of salinity, temperature, and pressure.

aint!

Density of seawater at the surface of the open sea ranges from
about 1.021 to 1.027 g/cm3. Generally about 1.025 g/cm3, density
increases with increasing salinity and continually decreases with
increasing temperature, provided the salinity is greater than 24.7 9/oo.
For lower salinities there is a temperature of maximum density as for
pure water. A small increase of density occurs with increase in pressure,
which is important because of high pressures in the deep ocean.

Seawater exhibits the unique properties of solutions, namely,
osmotic pressure, depression of freezing point, elevation of boiling
point, and reduction of vapor pressure. The observed freezing point
is -1.910°C for a salinity of 35 °/oo, the freezing point depression
being less than that computed on a basis of complete ionic dissociation
by a factor of 0.89. For all practical purposes, the latent heat of
evaporation may be taken as that of pure water.

Specific heat at constant pressure Cp is close to 1 cal/g°C and
increases with increasing pressure. The ratio of specific heat at
constant pressure to that at constant volume increases with temperature
and pressure. Seawater is compressible, and effects of adiabatic pro-
cesses, although small, must be taken into account when studying the
vertical distribution of temperature in the great depths of the oceans.
Adiabatic processes are those that take place without the flow of heat
into or out of the volume of the medium under consideration. When a
sample of water is brought in an insulated sampling bottle from an
initial depth (pressure) to the surface, a certain amount of thermo-
dynamic cooling occurs because of the change in pressure, irrespective
of the temperature of the local environment. The temperature that the
sample would have when brought from the initial depth to the surface
without heat flow is called the "potential temperature."

In the ocean, the important mechanical mixing property is eddy
viscosity, which is the effective viscosity occurring on a macroscopic
scale in large bodies of fluid. This concept is to be differentiated
from molecular viscosity, which is measured on a small sample of fluid
in the laboratory. Dynamic viscosity is defined as the ratio of
shearing stress to the resulting velocity gradient and expressed in
grams per centimeter per second. It is common in.fluid mechanics to
use kinematic viscosity, which is dynamic viscosity divided by density
and expressed in square centimeters per second. The molecular viscosity
for water is about 0.019 cm2/s at 0°C and approximately half this at
20°C; the eddy viscosity is substantially larger and may range from 2 to
7,500 cm2/s depending upon conditions. Surface tension controls the
capillary waves from which larger surface waves receive energy from the
wind. The surface tension value is about 76 dyn/cm at 0°C and 73 dyn/cm
ete ZO.

c. Optical and Radio Properties

Reduction of the intensity of optical radiation in the sea is
of major biological interest. In general, light is confined to the

172

upper few hundred feet of the ocean. Extinction of light is due to
both absorption and scattering, the former being greater at the red

end of the spectrum and the latter at the blue. Extinction is largely
a function of the amount of suspended material and is considerably
greater in coastal waters. The refractive index of seawater depends
strongly on salinity and is used to measure salinity (with appropriate
temperature corrections). Electrical conductivity rises rapidly with
an increase of salinity, and as a result electrical conductivity
measurements are used also for the determination of salinity. (Again,
there is a strong temperature dependence and a correction must be made.)
Conductivity, which is approximately 4 mho/m, limits the transmission
of radio waves. Figure 4-4 compares absorption of radio waves and

sound waves in seawater. It may be seen that sound is transmitted
better below about 100 MHz. As seawater has a relative permittivity

of 80 for radio waves, it is a conductor below 10 MHz and a dielectric
above 102 MHz. The skin depth 6 at which radio waves decay to l/e of
their surface field intensity value is given by 8 = 250/./f meters,
where f, the frequency, is in Hz. At 1 MHz, § = 0.25 m. Ten kW of
power in a plane wave at 16 kHz is attenuated to lwW at a depth of 23 m.
The speed of radio waves is reduced considerably by the conductivity,
being equal to 27f6. Thus, at 1 MHz, speed equals 1.5 x 10° m/s and is
5.0 x 104 m/s at 1 kHz.

d. Sound Speed

Of the many acoustic parameters associated with the ocean and
its boundaries, the most important are speed of sound and absorption
loss.

The speed of sound, C, in seawater is given by the formula
c2 = (dp/dp)s; where S denotes an adiabatic process. Sound speed is
related to the equation of state where p, the density, is given in
terms of pressure, temperature, and salinity. An equivalent form for
the speed of sound is c2 = 1/(p Ka) where Ka is the adiabatic compressi-
bility. The adiabatic compressibility is the reciprocal of adiabatic
bulk modulus. There are several expressions for the speed of sound in
terms of temperature, salinity,and pressure. The formula most widely
accepted is that of Wilson shown in figure 4-5.

At O°C, the speed of sound is 1449.3 m/s for a salinity of
35 °/oo at atmospheric pressure and increases by about 1.4 m/s per
1 °/oo increase in salinity, by 4.5 m/s per °C increase in temperature,
and by 0.016 m/s increase in depth of 1m (1 atmosphere ~10 m). These
variations with depth can result in marked acoustic refraction in the
ocean.

e. Sound Absorption

Absorption of sound in seawater increases rapidly with acoustic
frequency and is dependent on salinity, temperature, and pressure.

13

Absorption values are approximately 0.065 dB/km at 1 kHz and 1 dB/km
at 10 kHz (T = 4.0°C). The absorption coefficient is that component
of the attenuation coefficient that passes directly into heat in the
ocean.

Absorption of sound in pure water, computed from Rayleigh's
formula based on shear viscosity, is about one-third of the measured
absorption. The water molecule has a bent-triangle structure, with an
angle of 109° between the oxygen and two hydrogen atoms. Two modes
of packing exist for water molecules, ice and liquid, both of which
are always present in varying relative amounts in liquid water. The
density maximum at 4°C and atmospheric pressure for the liquid is due
to a relative maximum amount of the closely packed structure. Molecular
packing is pressure dependent, and in the presence of a compressional
sound wave, there is a relaxation effect causing the packing of some
molecular aggregates to shift back and forth between liquid and ice
structure. The extraction of acoustic energy in this process is the
pure water absorption.

In seawater, the additional absorption over that of pure water
is due to a pressure sensitive relaxation process at about 70 kHz,
arising from a shift between hydrated and nonhydrated magnesium sulphate
ions (fig. 4-6). Magnesium sulphate accounts for 5 percent of the salts
present in seawater. The effect of sodium chloride, accounting for
78 percent of seawater salinity, is to inhibit absorption. An
expression for the absorption of sound in seawater for frequencies
above 2.5 kHz as a function of temperature, salinity, and pressure is
given in figure 4-7.

A relaxation occurring at about 1 kHz increases seawater
absorption further, by a factor of 10 over the magnesium sulphate
absorption, at very low frequencies below 1 kHz. This relaxation
process is due to the boric acid-borate system in seawater. Thorp's
empirical expression for absorption of sound in seawater over the
entire frequency range is given in figure 4-8.

5. The Steady State of the Ocean

a. Heat Transport by Winds and Currents

Since the average temperature of Earth as a whole is not
changing appreciably, then the total amount of heat received from the
Sun must equal the amount lost by radiation back into space. However,
at any one time and place, Earth is either warming or cooling. The
intensity of solar radiation is largely a function of latitude and
season.

Figure 5-1 shows the radiation balance between insolation and
outgoing radiation as a function of latitude. The surplus in the

tropical regions is balanced by the deficit in polar regions. The

14

resulting latitudinal temperature distribution is shown to vary from
about 28°C at the equator to about -1.7°C at the poles. Much greater
differences in temperature between the tropics and poles would be
experienced, if it were not for the relatively free circulation of
atmosphere and oceans, which compose the working substance of a gigantic
heat exchanger. Large-scale heat transport from warmer to colder
regions is due to atmospheric winds and ocean currents.

Air, which acts as a major mechanism for redistributing heat,
is driven by unequal heating of the Earth's surface to rise at the
equator and sink at the poles. However, air cannot circulate uniformly
in a single convection cell, since Earth's rotation, through the
deflecting force of the Coriolis effect, prevents a direct north-south
path for the air. Earth's rotation and the distribution of land and
water cause a three-celled vertical circulation system from the equator
to a pole (fig. 5-2).

In the oceanic environment, moving air produces waves and
drives the major surface currents. Once surface waters are set in
motion, they create distributions of sea-surface temperatures which
in turn alter the winds that drive the currents. There is, however,
a mean steady-state condition.

A wind-driven current does not flow in exactly the same direction
as the wind, but is deflected by Earth's rotation (fig. 5-3). This
deflecting force, the Coriolis effect, is greater at higher latitudes
and is to the right of wind direction in the Northern Hemisphere and to
the left in the Southern Hemisphere. Between 10°N and 10°S, the
current usually sets downwind. In general, angular direction between
wind and the surface current varies from about 10° in shallow coastal
areas to as much as 45° in some areas of the open ocean. The angle
increases with the depth of current, and at certain depths the current
may flow in the opposite direction to that at the surface.

Ocean current patterns, consisting of huge gyres, are related
to wind systems of the world (fig. 5-4). Water motion at the surface,
modified by Coriolis effect, responds to: (1) density gradients
caused by latitudinal heating effects, and (2) surface winds, which
themselves are modified by the Earth's rotation and the distribution
of land and water masses. Position of the gyres shifts only slightly
with season, except in the northern part of the Indian Ocean and along
the China coast, where monsoons cause the currents to flow in opposite
directions in winter and summer.

Ocean currents may attain speeds of 5 or more knots as reached
by the Florida current. In middle latitudes, the strongest surface
currents rarely reach speeds above 2 knots. At depth, currents are
generally slow, usually less than 0.5 knot. However, there are sub-
merged high-velocity streams, such as the Cromwell Current in the
Pacific Ocean.

15

A slow, density-driven, vertical circulation occurs in the
ocean that leads to a stable density layering or stratification. Water
is heated in the tropics. This heated water decreases in density and
expands, spreads northward and southward along the surface, and gradu-
ally cools as it approaches higher latitudes. As warm surface water
is drained off from the tropics it is replaced by upwelling cold water
that originated in polar regions. The deepwater layer of all oceans
is derived from polar regions and has about the same characteristics
of salinity and temperature as the surface waters of those regions.
This happens because the water is uniformly cold in these regions and,
along with brine residue left behind in ice formation, produces the
heaviest waters in the ocean. This water sinks and spreads from one
pole to the other. The horizontal net drift in these deep thick layers
is very small, about 1 to 3 cm/s. This produces a low average tempera-
ture about 3.8°C for the whole ocean, and a conservative temperature—
salinity relationship.

b. Three-Layered Vertical Temperature Structure

Local convection is important in heat exchange near the surface.
Surface waters may be warmed down to 10 meters by solar radiation, and
may be cooled by evaporation, precipitation, and radiation. Both cooled
water and more saline water will be heavier than that beneath it and
will sink. These processes together with wind result in vertical mixing
and widespread formation of an isothermal surface layer.

The basic vertical thermal structure is essentially a three-
layered system (fig. 5-5): (1) a relatively warm thin surface layer
mixed by wind and convection processes, (2) a very deep mass of much
colder water in which the temperature decreases very slightly and
uniformly with depth, and (3) a transition layer, known as the main
thermocline. This structure is an inherently stable one since the
density, which is largely controlled by the temperature, increases
with depth (pycnocline). The warm surface laver is thickest in mid-
latitudes as is the thermocline or pycnocline. In latitudes higher than
about 50°N and S, no surface layer exists, and the entire water column is
the same temperature as the cold deep ocean water. Superimposed upon this
stable situation are marked diurnal and seasonal temperature variations
in the upper layers. Illustrating the three-layered ocean, figure 5-6
shows actual plots of temperature-depth curves for three different
latitudes for summer and winter.

Figure 5-7 indicates seasonal effects in the near-surface layer
due to wind and convective mixing. If the spring condition is a
moderately deep mixed layer, then surface heating and calm conditions
of summer will create a surface thermocline leaving an isothermal
section in the profile (fig. 5-7A). Winds then mix the near-surface
water creating a pair of mixed layers and thermoclines (fig. 5-7B).
When strong winds blow, one deep mixed layer develops (fig. 5-7C). In
autumn, surface cooling (fig. 5-7D) will induce convective mixing
(fig. 5-7E), and by winter the mixed layer is very deep (fig. 5-7F).

16

Figure 5-8 shows frequency of occurrence by percent of mixed layer
depths in the North Atlantic by season. In summer and fall they are mostly
shallower than 250 feet (76 m). In winter they are mostly deeper than

250 feet.

c. Salinity and Water Masses

Characteristics of ocean waters are controlled by surface
processes, for it is only at the sea surface that heat or freshwater
is added or removed in quantity. Thus, salinity and temperature of
midocean surface waters are determined by the distribution of evapora-
tion and precipitation processes as well as solar radiation. Surface
salinities reach maximum in belts centering at about latitudes 23°N and
S. Evaporation is maximum along the outer edges of the trade wind belt,
where the air is descending and thus relatively dry. More precipitation
than evaporation occurs in the tropics and at latitudes above 40°N and
S. Salinity in the surface waters of the ocean is proportional to the
excess of evaporation over precipitation. This is demonstrated in
figure 5-9.

Salinity effects on density are sizable, but the pycnocline is
controlled by development of the thermocline because of the relatively
large temperature range (-1.7°C to 30°C) of seawater compared with the
relatively small salinity range (33 to 37 °/oo).

In surface waters, the relationship between temperature and
salinity is highly variable because of surface heating and cooling,
evaporation, and precipitation. In the main thermocline and deepwater
layer, however, a direct correlation between temperature and salinity
exists. Knowing the temperature and water mass, one can predict the
salinity with almost as much accuracy as it can be measured by the
routine process (+0.02 °/oo).

The permanency of the oceans' gross vertical structure, main-
tained by vertical circulation, allows identification of water masses
from their vertical salinity and temperature distributions. If the
vertical distributions of temperature and salinity are plotted on a
graph with temperature and salinity as coordinates, and density excess
( Of ) as a parameter, then the resulting characteristic curve is
called a T-S curve. Figure 5-10 shows T-S curves for the Atlantic Ocean.
Figure 5-11 shows the water masses of the world oceans classified
from their T-S curves.

d. Sound Speed vs. Depth Structure

The density stratification of the ocean, with the three-layered
temperature model, leads to a typical sound speed vs. depth profile
such as seen in figure 5-12.

In the thermocline, temperature changes rapidly and dominates
the sound speed profile. In the mixed layers, the temperature profile

IL7/

is isothermal, and in the deep layer it changes slowly. In deep-layer
regions, the pressure effect on sound speed prevails. Direct salinity
effects on sound speed are usually small.

Figure 5-12 shows that when an isothermal mixed layer exists as
in winter, sound speed increases with depth from the surface at the
rate 0.016 m/s/m until the temperature effect of the thermocline takes
over. Beneath the thermocline, decrease of temperature with depth
becomes small, the pressure effect again takes over and sound speed
increases again close to the isothermal rate of 0.016 m/s/m. In
figure 5-12, another velocity minimum occurs followed by an increase
to the highest sound speeds in the ocean near the bottom. Total
variation in sound speed is of the order of about 4 percent depending
on water depth. This small variation along with the sound speed minima
has a decided effect on propagation of sound over long ranges.

Sound speed minima are axes of sound channels where rays are
trapped or concentrate, whereas the maxima cause a divergence among
rays. The sound-trapping region near the surface is called the surface
duct, whereas the sound-trapping region at depth is called the deep
sound channel or SOFAR (Sound Fixing and Ranging) channel. Figure 5-13
shows the latitudinal behavior of the deep sound channel axis depth in
the North Atlantic. The deep sound channels are deepest in midlatitudes.
An examination of the temperature vs. depth plots of figure 5-6 indicates
that this is caused by: (1) very steep thermoclines and shallow mixed
layers of tropical regions, and (2) near uniformity with depth of
temperature and salinity caused by convective mixing in high latitudes.

e. Modes of Sound Propagation in the Ocean

The typical sound speed profile of deep oceans results in four
major modes of sound propagation, depending on the depths of the sound
source and receiver. In figure 5-14A, the surface duct traps rays by
successive bounces off the surface and upward bending by the increasing
sound speed with depth. In figure 5-14B, rays from a source at the
deep sound channel axis are trapped by continued refraction due to the
positive gradients on either side of the channel axis. In figure 5-14C,
the convergence zone mode returns a ray bundle by total refraction to
the surface from great depths. Figure 5-14D shows the bottom and
surface bounce mode that prevails for steep rays, relatively shallow
water depths, or high surface temperatures.

6. Waves in the Ocean
a. The Sea Surface and Acoustic Effects
The state of the sea surface has important consequences for
ocean acoustics. When wind blows over the water, there is a transfer

of energy from wind to water that sets the sea surface in motion.
Surface waves are generated and a surface drift current is created.

18

About 20 percent of the water energy goes into surface drift currents
and the rest into surface waves. Ordinary wind-driven waves are
classified as capillary waves if surface tension controls the forma-
tion, or gravity waves if gravity controls wave structure.

The shape of deepwater waves depends to large extent on their
length. Gravity waves that are longer than about 1.73 cm are controlled
by gravitational and inertial forces. Wave shapes range from a nearly
sinusoidal form to a sharply peaked crest and rounded trough which can
be approximated by a trochoid (fig. 6-1). A trochoid is the curve
generated by a point on the radius of a circle as the circle rolls
uniformly on a line. Water particles execute small circular orbits in
response to the passing wave, the size of the orbits diminishing with
increasing depth below surface (fig. 6-2). Surface waves shorter than
1.73 cm in length are generally controlled by forces associated with
surface tension, causing the short capillary waves to have rounded
crests and v-shaped troughs. Figure 6-3 shows the shape of capillary
waves in terms of the amplitude-to-wavelength parameter (@/d).

Ocean waves are dispersive, since wave speed depends on wave-
length or frequency (fig. 6-4). Gravity waves show normal dispersion,
since the wave speed increases with increase in wavelength. Capillary
waves show anomalous dispersion, since their wave speed increases with
decreasing wavelength. The minimum wave speed for combined gravity
and capillary waves is 23 cm/s, corresponding to a wavelength of 1.73 cm.
The wave frequency » , is_related to wave number k, for gravity waves
through the expression: w“=gk where g is the gravitational constant.
For capillary waves the relationship is: w= Tk3/p, where: p is the
water density and 7 is the air-water surface tension coefficient,
equal to 73 dyn/cm at 20°C.

Wind effects on the sea surface are not simple, because turbu-
lence and viscosity in both sea and air introduce complex nonlinear
effects. Wind turbulence creates a moving pattern of minute fluctuations
in air pressure over water, which can generate the initial tiny ripples
that eventually become fully developed waves. It is thought that much
of the grip that wind has on water to produce waves and currents is
provided through those very small wavelets, since they are very numerous
and move slowly before the wind. Gravity waves speed up with growth
to the point where they can eventually move at the same speed as the
wind and can extract no further energy from it. Swell is the name
given to those gravity waves that move away from the generating area
faster than the wind.

Ordinary gravity waves normally observed at sea are composed of
two varieties: swell, which is a long and relatively symmetrical wave,
having a period in the order of 10 seconds, produced by wind and
storms at some distance from the point of observation, and sea, which
is the local wind-generated wave motion, usually of shorter period,
with unsymmetrical slopes and steep or white-capped crests depending
upon the wind speed and fetch. Fetch is the length of sea surface over
which the wind blows without appreciable change of direction.

19

For practical purposes, a sea-state scale based on visual
estimates relating wave height, wind speed, and sea surface appear-
ance has been quite useful. One, based on the Naval Oceanographic
Office code, is given in figure 6-5.

An important application of the sea-state code to underwater
sound is the Knudsen curves for average ambient noise in water shown
in figure 6-6. The curves fall off at the rate of about 5 dB/octave.

Another application of the sea-state code in terms of crest-to-
trough wave height occurs in the propagation of sound involving ray
paths that bounce off the sea surface. Propagation loss per bounce
(a3) is given as a function of the product of mean wave height, h(ft),
and acoustic frequency, f (kHz).

@s = 1.64 (fh) 1/2 ap/bounce

This relation is plotted in figure 6-7.

b. Energy Spectrum of the Wind Waves

Actual waves on the sea surface are irregular, aperiodic, and
short crested. The wave spectrum concept may be used to describe
distribution of energy among waves of different frequency or wave
number. Wave spectrum must be studied in terms of probabilistic models
and measured and analyzed by statistical techniques. A wave model is
used to describe fluctuation in wave height at any point in terms of a
statistically invariant gaussian function of time.

As the wind blows and as waves grow, turbulent variations and
varying amounts of internal eddy viscosity -- as well as the interaction
of one wave with another -- set an individual limit to the growth of
each individual component wave train in the developing sea. They do
this by initiating the energy-dissipating, breaking, or "white-cap"
process, whenever a momentarily high crest reaches an unstable configu-
ration. In classical theory, this occurs when the height-to-length
ratio of waves in a train reaches the critical value of 1/7 at which
point in a wave's development its crest is sharply peaked. Thus, the
total energy present in all waves on a developing sea progressively
distributes itself over a range of frequencies, each frequency charac-
terizing a particular wave train. As waves continue to grow and as new
trains continually develop, this range extends more and more to lower

frequency -- or longer period -- waves. In brief, a spectrum of ocean
waves is formed (fig. 6-8), in which -- for any given wind velocity
and for fully developed waves -- energy distribution over a narrow band

of wave frequencies from 0.05 to 0.3 Hz is distinctive.

The accepted expression for the energy spectrum of wind-driven
gravity waves is that of Pierson and Moskowitz (PM). It is a directional
spectrum and is given by:

20

S (0,7) = (ag?/o) [exp -8 (%/o) 1 F (o, 0%») ms
where: A =) 8h x 105° m2/s, 6 = 0574 and @ = 8/¥19.5> rad/sec

¥19.5 is the wind speed at 19.5 meters above the sea surface,
and F (o, 0*, v) is the beamwidth of the wave spectra depending on wave
angular frequency (), wind speed (v), and direction (@*) measured with
respect to direction of the wind and is between + 7/2. For the fully
risen or equilibrium sea, this spectrum falls off asw- 2, The exponential
term gives growth portion of the curve.

c. The Thermocline and Internal Waves

Fluctuations, variations, and waves in the temperature
structure of the thermocline occur quite frequently in the ocean and
can have important consequences for duct propagation. The sharp density
gradients of the thermocline give rise to gravity oscillations, called
internal waves (fig. 6-9). Oceanic internal waves generally have wave-
lengths, periods, and amplitudes much larger than surface waves. Internal
waves travel much slower than surface waves. These waves may have
periods measured in minutes and hours, wavelengths in kilometers, and
amplitudes of the order of tens of meters.

There are theoretical maximum and minimum frequencies for free
internal waves. The theoretical maximum frequency is known as the
Vaisala frequency. The theoretical lower limit is defined by the
Coriolis parameter. The Vaisala frequency N is expressed in terms of
density p , density gradient dp/dz, gravity g, and sound speed C, in
the medium as follows:

NhediPi 39

N?= 9 Uiomaa Ga)

The maximum frequency is a function of depth; it is largest in the
thermocline and least at depth of weakest density gradient. A contin-
uous monotonically decreasing power spectrum seems to exist between
minimum and maximum frequencies along with line frequencies of internal
waves.

Several different types of forces have been identified as
causing internal waves. Tidal and other forces that impel water
movement around land boundaries and topographic features can start
oscillations in the thermal structure. Strong winds may create
convection cells and eddies in the upper layers of the sea, and the
resulting circulation will cause current shears and waves on the
thermocline. Vertical oscillations in the thermocline can be induced
by air pressure changes and fluctuating winds at the sea surface.
Currents crossing sills or underwater mountain ranges may lead to
internal waves.

Power spectra of internal waves seem to show regions of distinct
frequency behavior (fig. 6-10). Power spectra show that the greatest

21

power density is in the low-frequency region. The power spectra
decrease monotonically with increase in frequency until a frequency

of about 0.02 cpm, at which point they fall off less rapidly with
frequency at about £-5/3 (k-l). At a frequency of about 0.8 cpm there
appears to be another transition zone corresponding to decay at a rate
Of £72) 20 (k=2).

d. Acoustic Propagation and Internal Waves

Horizontal thermoclines show a shadow zone beyond the limiting
ray in the duct. Within the duct, more or less acoustic energy is
trapped depending upon whether the limiting ray is steeper or shallower
than with an undisturbed thermocline. The effect of internal waves
on acoustic propagation is to form regions of convergence and divergence
of rays below the thermocline. In particular, figure 6-11 demonstrates
the effect on propagation in and below a surface duct. There is an
increase of energy leakage out of the duct bounded by a wavy thermocline.
In addition, there are alternating regions of high and low intensity
below it.

22

SURE 2-1. THE INTERIOR LAYERS OF THE EARTH. (ADAPTED FROM M.G. GROSS, OCEANOG-
RAPHY, CHARLES E. MERRILL, COLUMBUS, OHIO, 1967.)

CRUST

MANTLE

OUTER
CORE

INNER
CORE

LAYERS OF THE EARTH AND THEIR DENSITIES

DENSITY

g/cm3
LAYER DEPTH RANGE APPROXIMATE
km AVERAGE

CRUST

0-33

MANTLE 33-2900 4.5
OUTER CORE 2900-5100 11.4
INNER CORE 5100-6371

TOTAL EARTH

23

008 = 60h SZ. M é f e091 OPI .0ZE = .001

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27

ELEVATION (km)

10

oo

loa)

>

10

12

FIGURE 2-6. THE HYPSOGRAPHIC CURVE.

EARTH’S AREA IN 108 km?
1 2 3 4 5

% LAND AREA
20 40 60 80 100

MT. EVEREST 8.84 km

MEAN LAND
ELEVATION .84 km

SEA LEVEL

80 60 40 20 0
% OCEAN AREA
DEEPEST TRENCH 10.85 km

20 40 60 80 100
% EARTH’S AREA

28

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29

FIGURE 2-8. TYPICAL METHOD OF CONTINUOUS SEISMIC PROFILING, NORMAL INCIDENCE RE-
FLECTION ABOVE AND OBLIQUE REFLECTION BELOW, USED TO MEASURE SUB-
BOTTOM STRUCTURE. (ADAPTED FROM M.N. HILL, ED., THE SEA, VOL. 3, INTER-
SCIENCE, NEW YORK, 1963.)

SUFFICIENT DISTANCE TO APPROX.

SHIP . MINIMIZE SHIP’S NOISE 10 ft
UNDERWAY | ar (= 800 ft)
SPARK AND RECEIVER SEAR

HYDROPHONE

“nN

310 5 ft
TOWED ASTERN BELOW WATER SURFACER

WATER BOTTOM

Ye

i tall
pore ifid:
iiabii

\N'SUB-BOTTOM \SSOGEAK KKK REA XS AS NOX
\ . : aN > YX \ NS \ SS N Ny
REFLECTING HORIZONS = NOOO ox
ee en Yate Soe

HYDROPHONE MOVED AWAY FROM
T OR TOWARD SPARK SOURCE
a

FLOAT FLOAT

PREFERABLY
SHIP LYING TO

WATER BOTTOM

\ SUB-BOTTOM
REFLECTING HORIZONS

30

BOTTOM LOSS FOR “FAST” AND “SLOW” SEDIMENTS—RAYLEIGH MODEL

(ACOUSTIC FREQUENCY = 1 kHz).

LEGEND
Cp = SOUND SPEED IN BOTTOM
Cyw=SOUND SPEED IN WATER

Pp =DENSITY OF BOTTOM
P\yy=DENSITY OF WATER

SOUND “SLOW” SEDIMENT
Cp/Cw-0.9786

Pp /Pw = 1.47
POROSITY = 70% (SILTY CLAY)

~N

=

Br ee
Sp come

SOUND “FAST” SEDIMENT
Cp/Cw=1.196

Pg /Pw — 1.30
POROSITY = 40% (FINE SAND)

10 20 30 40 50 60 70 80
GRAZING ANGLE IN DEGREES

31

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33

FIGURE 3-3. A PLANKTON SAMPLE WITH REPRESENTATIVES OF THE MEROPLANKTON.
(NAVAL OCEANOGRAPHIC OFFICE)

ZOOPLANKTON

A DRIFTING COMMUNITY OF FEEBLY
SWIMMING ANIMALS FEEDING ON
PHYTOPLANKTON AND BEING EATEN,
IN TURN, BY THE NEKTON.

Illustrations by J. Recknagel
(NAVOCEANO)

LARVAL BOTTOM FAUNA, TEM-
S PORARILY ZOOPLANKTON, GO
mantis shramp!- acer BACK TO THE SEA FLOOR FOR

worm

THEIR ADULT LIFE.

34

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35

FIGURE 3-5. VARIATION OF EFFECTIVE SCATTERING AREA WITH WAVELENGTH FOR SMALL
SOLID PARTICLES AND BUBBLES. (NAVAL OCEANOGRAPHIC OFFICE, 1966)

EFFECTIVE SCATTERING AREA
ACTUAL AREA

CIRCUMFERENCE
WAVE LENGTH

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FIGURE 4-4. ABSORPTION OF WAVES IN SEAWATER.

ABSORPTION COEFFICIENT IN dB/m

107

10°

103

10° 102 104 10° 108 1019
FREQUENCY IN Hz

4]

FIGURE 4-5. SPEED OF SOUND IN SEAWATER (WILSON)

C= 1449.144+C++Cp+Cs+Cstp
C= 4.5721t— 4.4532 x 10—7t? — 2.6045 x 10—4t8
+-7:989 |< 105°
Cp = 1.60272 x 10—'p + 1.0268 x 10~5p?
+3.5216x 10—%p3— 3.3603 x 10—'?p4
Cs=1.39799 (S—35)-+ 1.69202 10" ($-35)?
Cstp = (S-35) (— 1.1244 10-7t+7.7711 x 10-7?
+7.7016x 10—5p— 1.2943 x 10 -7p?+ 3.1580 10 —8pt
+ 1.5790 x 10-%pt?)+p (— 1.8607 x 10-4t
+7.4812x 10—%t?+ 4.5283 x 10 —*t3)
+p? (—2.5294x 10—7t+ 1.8563 x 10~%t?)
+p? (—1.9646x 10—°).

WHERE: C_ IS SOUND SPEED IN m/s
t 1S TEMPERATURE IN °C (—4°C < T <30°C)
S IS SALINITY IN ‘eo (0 %o <S <37%o0)
p 1S PRESSURE IN kg/cm? OR ATMOSPHERES

(1kg/em? <P <1000kg/cm?)

J. ACOUST. SOC. AM. 32, 1357 (1960).

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FIGURE 5-1. RADIATION BALANCE IN THE NORTHERN HEMISPHERE AND AVERAGE OCEAN |

SURFACE TEMPERATURES. (ADAPTED FROM M. G. GROSS, OCEANOGRAPHY, SECOND EDITION,
CHARLES E. MERRILL PUBLISHING COMPANY, 1971, PAGE 73.)

CLIMATIC ZONES

TROPICS SUBTROPICS SUBPOLAR POLAR
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48

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49

DEPTH (KILOMETERS)

_
On

2.0

FIGURE 5-5. SIMPLE THREE-LAYERED OCEAN IN WINTER
SHOWING RELATIVE VOLUME IN EACH ZONE
(NATIONAL DEFENSE RESEARCH COMMITTEE
v. 6A, 1946).

LATITUDE

DEEP WATER

50

10°

FIGURE 5-6. REPRESENTATIVE WATER TEMPERATURE PROFILES BY SEASON AND LATITUDE.
(NAVAL OCEANOGRAPHIC OFFICE, 1966)

2000

3000

4000
[=
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a WINTER SUMMER
18 H
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a
9 MARCH 1935 ! 26 AUGUST 1958
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23 APRIL 1957 10 AUGUST 1958
(TEMPERATE ZONE) (TEMPERATE ZONE)
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23 NOVEMBER 1958 10 MAY 1957
(TROPICAL ZONE) (TROPICAL ZONE)
9000
10,000

WATER TEMPERATURE, °F

51

FIGURE 5-7. WIND AND CONVECTIVE MIXING.

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SURFACE HEATING MODERATE WINDS STRONG WINDS

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009

BIBLIOGRAPHY

Albers, V. M., Underwater acoustics handbook, Pennsylvania State
University, 1965.

Arrhenius, G., Pelagic sediments, The sea, vol. 3, The earth beneath
the sea, M. N. Hill, ed., Interscience, NY, 1963.

Charnock, H., A preliminary study of the directional spectrum of

short period internal waves, Proc. 2nd U.S. Navy Symp. Mil.
Oceanog., 175-178, 1965.

Crapper, G. D., An exact solution for progressive capillary waves of
arbitrary amplitude, J. Fluid Mech., 2, 532-540, 1957.

Duxbury, A. C., The Earth and its oceans, Addison-Wesley, 1971.

Earth Science Curriculum Project, Investigating the earth, Houghton-
Mifflin, Boston, 1967.

Fleming, R. H., General features of the ocean, Geological Society of
America Memoir, 67, 1957.

Forsyth, W. E., editor, Smithsonian physical tables, 9th rev.
(Washington, D.C.--Smithsonian Institution, 1964).

Frederick, Margaret Anne, An atlas of Secchi disc transparency
measurements and Forel scale color codes for the oceans of
the world, Thesis, U.S. Naval Postgraduate Sch., Sep 1970.

Gross, M. Grant., Oceanography, Second edition, Charles E. Merrill,
OV

Hardy, A. C., The open sea, Houghton Mifflin, Boston, 1964.

Hedgepeth, J. W., ed., Treatise on Marine Ecology and Paleoecology,
Ecology, 1, The Geological Society:of America Memoir, 67, Boulder,

Colorado, 1957.

Heezen, B. C., M. Tharp, and M. Ewing, The Floors of the Oceans, I. The
North Atlantic, The Geological Society of America, Special Paper 65,
New York, NY, 1959 (reprinted 1963).

Hersey, J. B., Continuous reflection profiling, The Sea, vol. 3, The
Earth Beneath the Sea, M. N. Hill, ed. Interscience, NY, 1963.

Houghton, H. G., On the annual heat balance of the northern hemi-

sphere, Journal of Meteorology, 11, 1954.

Knudsen, V. O., R. S. Alford, and J. W. Emling, Underwater ambient
noise, Journal of Marine Research, 8, 1948.

71

BIBLIOGRAPHY (con. )

LaFond, E. C., Internal waves, The sea, vol. 1, Physical Oceanography,
ed., M. N. Hill, Interscience, NY, 1962.

Lee, 0. S., Effect of an internal wave on sound in the ocean, J.
Acoust. Soc. of Amer., 33, 1961.

National Defense Research Committee, The Application of oceanography

to subsurface warfare, vol. 6A, Washington, D.C., ed. C. O'D Iselin,
1946.

Naval Oceanographic Office, Handbook of oceanographic tables (U),
U.S. Naval Oceanographic Office, SP-68, Washington, D.C., 1966.

Naval Oceanographic Office, Oceanography and underwater sound for naval
applications (U), U.S. Naval Oceanographic Office, SP-84,
Washington, D.C., 1966.

Office of the Oceanographer of the Navy, The ocean science program of
the U.S. Navy, accomplishments and prospects, Alexandria, Va., 1970.

Phillips, 0. M., The Dynamics of the upper ocean, Cambridge University
Press (1966), 261 p.

Pierson, W. J., and L. Moskowitz, A proposed spectral form for fully
developed wind seas based on the similarity theory of S. A.
Kitaigorodskii, J. Geophys. Res., 69, 5181-5203, 1964.

Reid, J. L., Observations of internal tides in October 1950, Trans.

Amer. Geophys. Un., 37, 278-286, 1956.

Revelle, R. and H. E. Suess, Gases, The sea, vol. 1, Physical oceanography,
M. N. Hill, ed., Interscience, NY, 1962.

Schulkin, M. and H. W. Marsh, "Absorption of sound in sea water,"
Ji bitte RE se 25199 5-500 CIOS).

Schulkin, M., The propagation of sound in imperfect ocean surface
duct, USL Report No. 1013, 22 April 1969, Naval Underwater
Systems Center, New London, Conn.

Strahler, A. N., Physical geography, Wiley, NY, 1960.

Sverdrup, H. U., M. W. Johnson, and R. H. Fleming, The oceans--their

physics, chemistry, and general biology, Prentice-Hall, 1948.

WD

BIBLIOGRAPHY (con. )

Thorp, W. H., "Analytic description of the low-frequency attenuation
coefficient," J. Acoust. Soc. Am., 42, 270(L), 1967.

von Arx, W. S., Introduction to physical oceanography, Addison-Wesley,
Reading, Mass. (1962). :

Wilson, W. D., "Equation for the speed of sound in sea water,"

Jour.
Acoust. Soc. Am., 32, 1357, Oct 1960.

13

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NOO RP~-1
TITLE (and Subtitle) 5. TYPE OF REPORT & PERIOD COVERED

BASIC ACOUSTIC OCEANOGRAPHY Final

6. PERFORMING ORG. REPORT NUMBER

AU THOR(s) 8. CONTRACT OR GRANT NUMBER(®)

Morris Schulkin

PERFORMING ORGANIZATION NAME AND ADDRESS a Naaeaae TASK

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o 0 Apri
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13. NUMBER OF PAGES
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SUPPLEMENTARY NOTES

KEY WORDS (Continue on reverse side if necessary and identify by block number)
submarine geology sound speed and sound absorption

bottom acoustic loss steady state of the ocean
marine biology sound speed vs. depth structure
deep sound scattering layer modesof sound propagation

ABSTRACT (Continue on reverse side if necessary and identify by block S GnbeD

A brief introduction is given of oceanography for acousticians. The main
topics and related acoustical examples are: submarine geology (bottom
acoustic loss); marine biology (the deep sound scattering layer); chemical
and physical properties of seawater (sound speed and sound absorption) ;
the steady state of the ocean (sound speed vs depth structure and modes of
sound propagation in the ocean); and waves in the ocean (sea surface loss
and sea surface noise, and acoustic propagation and internal waves).

DD , Fees 1473 EDITION OF 1 NOV 65 1S OBSOLETE UNCLASSIFIED

S/N 0102-014- 6601 | ee
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UNCLASSIFIED

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waves in the ocean
sea surface loss

sea surface noise
acoustic propagation
internal waves

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ree spt dp OR

se

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