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TECHNICAL NOTE 9-SAIL-1

Proceedings of the

Sea-Air Interaction Conference
Tallahassee,Florida ae
‘February 23-25, 1965 <o

"TECHNICAL NOTE 9

“SEA-AIR INTERACTION LABORATORY
REPORT NO.1

_+ WASHINGTON,D.C.
August 20, 1965

WEATHER BUREAU TECHNICAL NOTES

SEA-AIR INTERACTION LABORATORY REPORTS

The Sea-Air Interaction Laboratory of the Environmental Science
Services Administration is responsible for the main research effort in
problems concerning the physical aspects of exchange processes between the
ocean and the atmosphere. The laboratory is also charged with the respon-

sibility of coordinating the federal effort in air-sea interaction research
and the development of a federal research program.

The laboratory's program encompasses research in an in-house basis such
as storm surge research and the data collection program at sea.

Other
projects are accomplished on a contract basis by universities and other
private agencies.

Reports by the laboratory staff, contractors, and cooperators will be
preprinted in this series to allow immediate distribution of the information
among the workers and other interested organizations. Since these reports
may not be in completely polished form and are for limited reproduction and
distribution they do not constitute formal scientific publication.

Reference
n this series puculd suionrd it as a preprinted report. Formal

made later in appropriate

Clearinghouse for Federal
. Department of Commerce, Sills
rginia 22151.

BL/

WMO

301 00

U.S. DEPARTMENT OF COMMERCE e John T. Connor, Secretary

ENVIRONMENTAL SCIENCE SERVICES ADMINISTRATION

Robert M. White, Administrator

Weather Bureau

TECHNICAL NOTE 9-SAIL-1

Proceedings of the Sea-Air Interaction Conference

Tallahassee, Florida,February 23-25, 1965

SPONSORED JOINTLY BY

FLORIDA STATE UNIVERSITY
DEPARTMENT OF METEOROLOGY AND OCEANOGRAPHY INSTITUTE,

AND

ENVIRONMENTAL SCIENCE SERVICES ADMINISTRATION,
SEA-AIR INTERACTION LABORATORY

WASHINGTON, D.C.
August 20, 1965

telat

FOREWORD

The interactions between the atmosphere and the oceans have increasingly
attracted the attention of meteorologists and oceanographers in recent years.
It is generally agreed that progress in our understanding of, and ability to
predict the behavior of the atmosphere and the oceans depends on a knowledge
of the exchanges of matter, momentum,and energy between the two fluid systems.
At the present time, adequate quantitative information on these exchange
processes at the air-sea interface is lacking.

It tas come to be recognized that a strong interdiciplinary research
program in meteorology and oceanography is required for a more satisfactory
comprehension of the interactions between the two parts of the coupled air-
sea system. In several universities and other research institutions, joint
programs of meteorological and oceanographic research are being conducted
with emphasis on exchange phenomena at the sea surface, and in the atmospheric
and oceanic boundary layers. The National Academy of Sciences has identified
air-sea interaction research as one of the scientific areas requiring increased
support, and the Federal Government has responded by establishing special
programs for this purpose.

The Sea-Air Interaction Conference held in Tallahassee in February 23-25,
1965, was suggested and arranged by Prof. Charles L. Jordan, Chairman of the
Department of Meteorology at Florida State University, in cooperation with
‘the Oceanography Institute of the University. The conference was co-sponsored
by the Sea-Air Interaction Laboratory (SAIL) of the Department of Commerce
(a joint laboratory of the U.S. Weather Bureau and the U.S. Coast and Geodetic
Survey). Mr. Feodor Ostapoff, acting director of SAIL, served as editor of
the Proceedings and arranged for its publication.

In the congenial and informal atmosphere provided by the conference hosts
at Florida State University, meteorologists and oceanographers from many
universities, research institutions and government agencies exchanged views,
reviewed programs and problems, and laid the groundwork for future cooperation
in air-sea interaction studies.

The participation of all the conferees in the work of the conference, and
the promptness with which they provided their papers to the editor for
publication in these Proceedings, are gratefully acknowledged.

Jerome Spar
Director of Meteorological Research,
U. S. Weather Bureau

TABLE OF CONTENTS

POTS WON Gis otetete svsvsrevcvetatel ctedel oneieilssilel cieNebeienensuenevre’ el exe: ciel: eifet alisiel sieriavelisiaveleielevemal cxeworekeveeyeclolcls

"Air Sea Exchange as a Factor in Synoptic-Scale Meteorology
algo: Whl@lsbiS inexeslcles\. oie dieiaoMme Bisebos oSaadadgcoodacsddos00cdsbs00 6b

"The Three=Dimensional Ocean Circulation Driven by Density
Gradients in an Enclosed Basin, " by Dr. Kirk Bryan.............. 17

"On the Present State of Knowledge in Air-Sea Boundary Layer
Problems.; "" spy Drip MisgiUie ROM. one roveier lop stere ole\elevevererenes a oer cusveisrete oletevetereie? nie

"A Survey of the Role of Sea=-Air Interaction in Tropical
Meteorollosy ik iby sDeedOannem\alicnsl SamosOneen eee OD)

"Sensible and Latent Heat Exchange in Low Latitude Synoptic-Scale
Syscems,. by Dir MalchaeieuGarsitangs «1.1 -1-lejaleielarcleleatereichetels cieveletotereyersterl Os)

"Intensity of Hurricanes in Relation to Sea Surface Energy
Pxchansel.” by Drvinew Perino osrverere oleceverelelscevelcneretorerietersterelsrerenereereted ile

"The Gulf of Mexico after HILDA (Preliminary Results)," by Dr. Dale
TD TS 0} 01-9 ae RORY C15 CRC CLIT RAR ie chs he IN cs Osteo IOS

"Evidence of Surface Cooling Due to Typhoons," by Dr. C. L. Jordan....185

"The Modification of Water Temperatures by Hurricane CARLA," by
Drs. Robert E. Stevenson and Reed S. Armstrong........0cesce+oee LG

“ion the Low Level Thermal Stratification of the Monsoon Air over

the Arabian Sea and its Connection to the Water Temperature
PASC, “leony Wies dess Ao Collie soccoovccoco cnc don cvanocooN 00000000 LOS)

J Low-Level Jet Produced by Air, Sea, and Land Interactions," by
DPS AMGwS WMHs: IBUTICST sarees ee evereievelerel cuelevenevele ove ciel eievensterersnerele eeleloterevers chee lO

"U. S. Fleet Numerical Weather Facility Activities Relating to
Sea-Air Interactions on a Synoptic Seale," by Comdr. W. E.
HETUNID Ae tayav eval aces erat onan welsova’ a}.e)reive Verieliare' ete Nel/oueVel ellay epenetevey etval’eveleileteleveWenavecelioleverorsceteveierer susie Oo,

"Synoptic Scale Heat Exchange and its Relations to Weather," by
DeLay OMe Vial Sigllicrstereteloleroeleneletencdelarelclenensleraletetel ciel slichercberslichevelerebenelsleneneveneieil 4

"Laboratory Studies of Wind Action on Water Standing in a Channel,"
bye Drs George Memiidy, rand Mewdimmrieab orm tise de siiieilerelscricii tector 2ob

"On the Instability of Ekman seca yee by Dr. ee K.

Tanya see See anes ote j Uawisctnae ;

"Federal Research Programs in Air-Sea Interaction," by Feodor

Ostapoff....... Sod00CdODOGaOE Sadd0d000Gd0000b0006

Appendix I: Conference Program.....-e-cocccceserecrrereercerrecsce

Appendix II: Conference Participants.....--ccsercceres

327

Bem
341
344

ATR-SEA EXCHANGE AS A FACTOR
IN SYNOPTIC-SCALE METEOROLOGY IN

MIDDLE LATITUDES

Jerome Spar
U. S. Weather Bureau, Washington, D.C.

INTRODUCTION

The subject of this conference is not a new one. Oceanographers and
meteorologists have been concerned with exchanges across the air-sea
boundary for many years and they do not need to be convinced of the
importance of these exchange processes.

What we have assembled to discuss here, I presume are the unsolved
problems of air-sea exchange, and the possibilities for solving these
problems through new measurement programs, theoretical investigations, and
computations. In particular, we need to concer ourselves now with improve-
ments in our quantitative information about the exchange processes, and
with the establishment of useful relationships between the microphysics
of the exchange processes and the larger scale atmospheric and oceanic
parameters. [

One of the important subjects for this conference should be “interaction”
in its full sense. Exchanges between the air and sea are generally viewed
either from above or below =-- rarely from both viewpoints simultaneously.
Most of us tend to interpret "interaction" parochially as the transfer of
property to or from our fluid (whether air or water) without regard to the
subsequent effect on the other side of the boundary. I suppose we all
cherish the long-term objective of treating the air-sea system as an
integrated problem. At the moment, however, few of us are prepared to do
very much about the real interaction problem. But perhaps we may be able
soon to make a crude assessment of the relative importance of the reaction
on one fluid of changes induced by that fluid in the other. How tightly
are the air and sea linked? Are the interactions strong or weak? Do
differences in the response times of the two fluids cause them to behave
as if, for practical purposes, they do not interact, but only influence
each other? ay

All oceanographers recognize that the atmosphere exerts an important
influence on the oceans. Clouds affect the distribution of insolation.
Rain falls into the sea. Evaporation varies with wind and humidity of the
air. The wind stress drives currents, generates waves, produces upwelling,
stirs the water, and creates spray. Atmospheric gases enter the sea, and
heat is exchanged between the air and the sea.

Every meteorologist knows that the oceans have a profound influence on
the atmosphere. Salt particles enter the air from the sea. Sea water
evaporates into the air. Through surface wind stress, the atmosphere loses
kinetic energy to the sea. And heat is exchanged between the sea and the
air.

We have come to recognize that through these exchanges of matter and
energy, a complex interaction takes place between the two fluid systems of
the earth. Thus the wind stress on the sea surface may change the sea
surface temperature distribution, which alters the heat transfer from the

sea to the air; and this in turn may change the atmospheric circulation,
thus altering the wind stress.

Nevertheless, meteorologists have found it convenient -- when they
have considered the oceans at all -- to think of the oceans and other large
water bodies as inert systems with inifinite heat capacity. In this way we
have been able to ignore “interactions” between sea and air entirely, and
have treated the sea surface as fixed in time, and unaffected by the
atmosphere.

To allow for the obviously non-negligible annual variation in sea
surface temperatures, we may use climatological mean monthly sea surface
temperatures, rather than mean annuals, and in some cases shorter time
averages. But, for the most part, the meteorologist has taken the ocean
state as "given" and has not attempted to predict it.

In extended and long-range weather forecasting and to a lesser extent
in short-range forecasting, some attempts have been made to employ the water
temperature anomalies as meteorological predictors. The underlying principle
in these efforts, however, is again the relative persistence of oceanographic
features -- in this case sea temperature anomalies -- compared with the var-
jability of the atmosphere. A one-way transfer, from sea to air, without
interaction, is implicit in these applications.

Of course, meteorologists are fully aware of the fact that the sea
responds to the atmosphere. But we have generally been unable to incorporate
oceanographic predictions in our work, either qualitatively or quantitatively.
In fact we have made hardly any progress even with the relatively simpler
one-way transfer problem.

We have come closest to dealing with a true air-sea interaction in
practical meteorology in connection with the hurricane problem. Several
investigators using synoptic ocean data, have presented interesting, albeit
inconclusive evidence pointing to the possibility that hurricanes may tend
to form over anomalously warm water, move along warm water anomaly "channels,"
avoid cold water, or dissipate over cold water. At the same time, these and
other studies have shown that the hurricane -- whether through upwelling,
stirring, evaporation or radiative-convective heat transfer -- cools the sea
in its immediate vicinity, and appears to leave a cold water wake behind the
storm. So large is this effect of the storm on the sea, that it is obviously
unrealistic to assume a fixed sea temperature anomaly field in hurricane
prediction.

To the oceanographer, looking at the underside of the air-sea boundary,
oceanographic prediction depends almost entirely on accurate prediction of
the atmosphere. Current anomalies, waves, and temperature anomalies -- as
well as salinity, oxygen and other oceanic anomalies -- are ultimately
meteorological in origin. The oceanographic prediction begins with a weather
map. Because of the time lag between storm development and wave generation,
plus the travel time of waves, a useful oceanographic prediction. can (like a

WwW

useful river forecast) often be made from the current weather map. But in
general the ocean forecaster requires prognostic weather maps. Abrupt and
unpredicted deepending or filling of a cyclone near a continental coast -- 2
phenomenon that could conceivably be caused by a wind-generated oceanic
anomaly -- can produce a disastrous failure in the prediction of sea state
and coastal wave action.

The interdependence of meteorological and oceanographic predictions on
alltime scales is too obvious to belabor further except to note that while
the atmospheric influence on the sea is reasonably well known, even quan-
titatively, the converse, i.e., the influence of the sea on the atmosphere,
is not.

Given the wind stress at the sea surface, we can at least compute the
steady state Ekman current as a function of depth. We can even generate
from the wind a reasonably realistic wave spectrum. True we do not know too
much about the depth to which the sea is mixed by surface wind action, and
we cannot predict temperature anomalies caused by upwelling with satisfactory
success -- but these problems appear to be amenable to solution.

The meteorological problem, on the other hand, seems to be much more
difficult. The effect of the air on the sea is direct and measurable. The
converse is neither direct nor measurable. The transfers across the sea
surface that are likely to have an important influence on the atmosphere
are of salt, water vapor, kinetic energy, momentum, and heat. None of these
transfers is easily measured, especially in high winds. But even if we
could measure them, we still have the problem of determining quantitatively
the effects of these transfers on the atmosphere.

Does anomalously high storm activity at sea measurably increase the
salt particle population of the atmosphere? How is the salt distributed?
What is its residence time in the atmosphere? How does it affect the subse-
quent precipitation distribution? What is the nature of the dynamic feedback
to the atmosphere of the latent heat thus released?

Similar questions may be asked regarding evaporative transfer, and again
the questions are not easily answered. We note that in an east coastal
cyclone (see, e.g., Petterssen,et al. 1962) most of the evaporative flux occurs
in the cold air to the rear of the surface cyclone. What fraction of the
transferred latent heat is realized locally in the cold air by cumulus
development and showers, and what fraction is realized far from the source,
perhaps in another system altogether? And what role does this energy play
in the cyclone development?

The sensible heat transfer from the warm sea to a cold air mass (the
major energy transfer according to Manabe, 1957) presents similar problems.
What do we really know about the dynamical effects on the atmosphere of this
sea-air heat transfer?

In the following I will try to review briefly some efforts that have
been made to evaluate the quantitative effects of sea-air heat and vapor
transfers on atmospheric circulation systems, largely in the context of the
weather prediction problem. (I will not attempt to deal with the important
salt transfer problem at all.)

QUANTITATIVE ESTIMATES OF SEA-ATIR TRANSFER
AND ITS EFFECT ON CIRCULATION

Investigations of this problem have almost universally ignored inter-
action effects, (i.e., thermodynamic feedback to the oceans), and it appears
very likely that such feedbacks may indeed be second order effects. The
ocean is thus taken as time-invariant.

Three approaches to the problem have been tried.

(1) Direct computations of the sea-air heat transfer (a) from the
transfer equations, following the Eulerian approach of Sverdrup (e.g.,Jacobs,
Petterssen, et al., ke), (b) empirically from trajectory (Langrangian,
air mass modification) studies (e.g., Burke, Craddock, Spar), and (c) from
line integral computations of ehergy flux divergence (Manabe).

(2) Indirect evaluation of Sgeseels heat transfers from a study of the
errors generated by adiabatic NWP’ models (e.g., Winston, Martin, Petterssen,
et al., Pyke).

(3) NWP computations with diabatic models, including sea-air heat

transfer (Bushby and Hinds, Reed, Spar).

A. Direct Transfer Computations

Direct computations of sea-air energy transfer do not really tell
us what role the transfer plays in the generation of circulation. WNeverthe-
less, the results are illuminating, and potentially useful for prediction.

To apply the Sverdrup-type transfer equations to the computation of
sensible or latent heat transfer, as Petterssen, et al. (1962) and Pyke
(1965) have done for diagnostic purposes, and others -- Bushy and Hinds (1955),
Reed (1958) and Spar (1962), for example -- have done in prognostic exercises,
one must have some knowledge of the transfer coefficients, and indeed one
really needs to know the correct functional form of, the transfer relation.
While it is convenient to assume a linear relation between the transfer rate
and the air-sea temperature or vapor pressure difference, and a linear
dependence on wind speed, as well, these assumptions are really not strongly
supported by data.

i/Numerical weather prediction

Empirical studies of two kinds have been conducted which could shed
some light on the problems of the functional form of the transfer relation
and the values of the transfer coefficients. However, these studies have
not been carried far enough to solve the general problem, and the data from
these studies have generally been used only te provide limited and immediate
practical answers for some special needs.

The one general result which does appear to emerge from these studies is
that we can assume zero sensible heat transfer in the stable case, i.e. where
warm air passes over cold water, and probably zero latent heat transfer as
well, as far as large scale dynamical effects are concerned. Obviously, the
well-known modifications of the shallow surface layer are important for
weather prediction; poleward moving surface air cools, and fog and stratus
do form. But while these results must be included in the complete weather
prediction computation, the total energy transfer involved is small, the
effect does not penetrate very high, and its dynamical consequences are
probably negligible.

The data from Burke's (1945) early experiment -- carried out at
Sverdrup's suggestion -- are unfortunately not presented in a form that
permits one to relate the sea-air energy transfer parametrically to macro-
scale variables. Craddock's (1951) data are somewhat more useful in this
regard. Several years ago (Spar, 1962) I tried to use the Lagrangian
trajectory technique, as Burke and Carddock had done earlier in their studies
of air mass modification, to evaluate the transfer coefficients for sensible
heat and water vapor. The results, based on 238 12-hour trajectories off the
east coast of the United States were the following:

In the case of "effective heat flux," (i.e. the sensible heat flux plus
radiative heating plus that latent heat released locally in the cold air by
cumulus formation and showers) the formula

i 0 Wa (BR. ow.) (1)

(H in ly day Se V. the average "surface" wind along the 12-hour trajectory
IM Tsecugs, Le and T. the average sea surface and "surface" air temperatures
in degree C) gave "satisfactory" results in the sense that the correlation
between the left and right hand sides of the formula was about 0.6.

The transfer coefficient above may be compared (although the comparison
is not strictly valid) with some others (see Table I).

Table 1. Sensible (and effective) heat transfer coefficients,
Ks, from various sources.
Ly-1

Dimensions: ly day~! (m sec™ degree La:

Ss Source
IOS Spar (1962)
3.6 Malkus (1962)
6.8 Jacobs (1942)

For water vapor the effort to evaluate a transfer coefficient from the
238 trajectories was less successful. The linear relation,

Ib Ska Vil(Gly = Gly) io (2)

(L, the latent heat transfer rate in ly day”? Vo the surface wind in
m sec, Gs and qo the dimensionless specific humidity at the sea surface
and in the "surface" air) could not be verified by the data because of the

large scatter (low correlation). A value of was determined, nonetheless.
from mean values of Vo (ds - do) and L. Table 2 shows this (dubious) value

of Ke together with some others.

Table 2. Latent heat transfer coefficients, K,,
from various sources.

Dimensions: ly dane (m seem) ve

Ke (x 103) Source

5.5 Spar (1962)

9.9 Marciano and Harbeck (1952)

8.5 Manabe (1958) (average over
all speeds) 1/

8.6 Malkus (1962) 7

Th 5 Jacobs (1951)

Manabe (1958) has applied the line integral method for computing the
horizontal flux divergence of latent and sensible heat to the Japan Sea with
very satisfactory results. Unfortunately, Manabe did not use his data to
test the parametric transfer formula for sensible heat as he did for evapora-
tion. This task remains to be done.

B. Indirect Computations of Energy Transfer

The errors in adiabatic numerical prediction models are in part due to
sea-air energy transfers, although other factors, notably condensation, may

17 (Manabe's results show a change in K, from 6.0 at low speeds - 4
6 m sec™~ - and a smooth surface, to” 11. at higher speeds -8 m sec -
and presumably a rougher surface.)

be even more significant. Studies of these diabatic errors leave little
doubt about the fact that they may be, on occasion, large and important,
and justify the inclusion of diabatic processes in NWP.

Winston's (1955) study of the February 1950 cyclogenesis in the Gulf
of Alaska, recently re-examined and extended by Pyke (1965), is an early
example of the efforts to evaluate the sea-air transfer from prediction
errors. A more comprehensive study of NWP errors by Martin (1962) demon-
strated even more clearly the probably importance of the sea as an energy
source for the atmosphere. Martin's computations were in remarkably good
agreement with the independent computations of Manabe (1958) for the winter
1954-55 cold outbreak over the Japan Sea (with transfers of more than
1400 ly day~1), and also with computations by Petterssen for the North
Atlantic.

Petterssen, Bradbury, and Pedersen (1962) in a diagnostic study of
eyclone development over the North Atlantic Ocean have attempted to develop
the classical Norwegian cyclone models into a more complete dynamical model
by computing the energy transfers from the sea to the air. The results have
been somewhat disappointing. In the first place, the computed heat transfers
show, as expected, the major heat transfer in the cold air mass to the rear
of the cyclone -- but with no physical account of how (or if) this energy
contributes to the cyclone development. Secondly, the paper essentially
bypasses the question of how the heat source affects the circulation, because
only the thickness (i.e., temperature) tendency is computed. Inclusion of
the heat source term in the tendency calculation improves the predicted
thickness tendency, as expected, but unfortunately, tells us nothing about
the effect on the 500-mb circulation.

C. Diabatic Prediction Models

The effects of sea-air heat transfer on atmospheric circulation systems
are so complex that it appears likely that nothing less than time integration
of the complete diabatic system of equations can really tell us much about
these phenomena. Such experiments have been attempted with models of vary-
ing degrees of complexity. Bushby and Hinds (1955) were probably the first
to incorporate sea-air heat transfer in a numerical weather prediction model,
followed a few years later by Reed (1958), who employed Fjgrtoft's graphical

(Lagrangian) method.

My own experiments (Spar, 1962), including heat of condensation as well
as sea-air heat transfer, employed a somewhat less constrained model, but
still only a two-level (vertically integrated) baroclinic model --
geographically limited and geostrophic. Despite these constraints, the
experimental results may be of some interest.

In the computations with my prediction model, I have used the empirical
transfer formulas (equations (1) and (2)) to compute (effective) sensible

heat transfer and evaporation. Input data included 1000 and 500 mb geo-
potential heights, integrated specific humidity (precipitable water), and
Ocean surface temperatures (time invariant). Output consisted of 1000 and
500 mb geopotential, precipitable water, cumulative precipitation, and
vertically - averaged vertical motion. The forecasts shown are for 12 hours
(computed in one-hour time steps).

The figures (Spar, et al., 1961) show the initial state, 12-hour
forecasts, and verification maps for a case of rapid North American east-
coastal cyclogenesis beginning at 0300 GMT 10 February 1957, and permit us
to compare forecasts made with and without surface heat and vapor transfer.

The initial conditions for the forecast experiment are displayed in
the three maps of Figure 1 which show the 500 mb geopotential and temperature,
1000 mb geopotential and 1000-500 mb geopotential thickmess, and the
vertically-averaged specific humidity.

Figure 2 illustrates the twelve-hour 500 mb forecasts computed (A)
with a barotropic model, (B) with a baroclinic model including heat of
condensation but no heat or vapor flux from the sea ("No Flux"), and (¢)
with a baroclinic model including sea-air fluxes as well as latent heat
("complete"). The verification contours are shown as dashed curves in
Figure 2(A). It is noteworthy that the baroclinic model predicted the marked
deepending that was missed by the barotropic forecast, but that the inclusion
of the sea-air fluxes had no significant influence on the twelve hour
forecasts.

The forecast 1000 mb maps, together with the predicted 1000-500 mb
thickness patterms are show in Figure 3 for the "No Flux" (A) and "complete"
(B) models. Also showm is the verification map (C) for the 1000 mb level
(solid curves) and the thickness (dashed curves). Again the forecasts made
with and without sea-air fluxes do not differ significantly in this short
time interval.

The conclusions drawn from these crude experiments were:

1. During cyclogenesis the diabatic effects (and this includes also
the latent heat release) were second order effects compared with baroclinic
effects. This conclusion, i.e., that potential to kinetic energy conversion
is the dominant energy transformation in cyclogenesis, is in agreement with
the results of many other investigators.

2. In 12 hours the dynamical effects of sea-air heat and vapor
transfer were of little significance. It might be expected that over a
longer period the air-sea effects might be more important. But it is
noteworthy that the 12-hour period selected was the one when the cold out-
break over the water was strongest, and the heat transfer at its peak.

ali

(Cc)

Figure 1. Initial conditions, 0300 GMT 10 February 1957. (A) 500 mb
geopotential height (solid curves, labeled in hundreds of feet) and
temperature (dashed curves, labeled in °C); (B) 1000 mb geopotential height
(solid curves labeled in hundreds of feet) and 1000-500 mb geopotential
thickness (dashed curves labeled in hundreds of feet), (C) mean specific
humidity (parts per hundred thousand). See text page 10.

Figure 2. Twelve-hour 500 mb forecasts for 1500 GMT 10 February 1957.
(A) Barotropic. Solid lines are the predicted contours; dashed lines
are the observed contours. (B) "No Flux." (C) "Complete." Contours
are drawn for an interval of 200 geopotential feet. See text page 10.

Figure 3. 1000 mb contours (solid lines) and 1000-500 mb thickness
contours (dashed lines) drawn for interval of 200 geopotential feet.
1500 GMT, 10 February 1957. (A) "No Flux," and (B) "Complete" 12-hour
forecasts. (C) Verification (observed) map. See text page 10.

Si Sea-air transfer did not affect the large scale vertical motion
in the system significantly in 12 hours.

h, The sea-air transfer effects were shallow, failing to penetrate up
through the 500 mb level.

These rather negative conclusions regarding the role of sea-air transfer
in the dynamics of coastal cyclogenesis must be viewed warily. The model
is constrained (notably in regard to the static stability); the experiments
were few in number; the period of integration was short.

Nevertheless the conclusions are not greatly at variance with those of
other investigators. Prof. Yale Mintz recently wrote the following to me,
in reply to a request for his views on the subject:

"I am sure this heating plays an important role in determin-
ing the temperature, wind and pressure fields, but in some
complicated way affecting more than just the cyclone scale of
motion.".... "If I have to guess at an answer, I would say
that the heat transfer from the sea affects the baroclinicity
of the air and hence the subsequent cyclogenesis; but that a
cyclone already in the developing state is, itself, relatively
little affected by the heat transferred to it. But that is
only a guess."

I am inclined to believe that this is a rather shrewd guess.

Jacobs, W.C.,

Petterssen, S.; D. L. Bradbury,
and K. Pedersen,

Pyke, Charles BR.

Burke, Cletus J.

Craddock, J. M.

Spar, J.

Sie, Ho8 dio Wo Caicicakair, dheos
L. A. Cohen

Manabe, S.

REFERENCES

i9he:

1951:

1962:

1965:

1945;

195i:

1962:

1961:

US /s

On the energy exchange between
sea and atmosphere. J. Mar.

Res., 5» 37-66.

The energy exchange between sea
and atmosphere and some of its
consequences. Bull. Scripps Inst.
of Ocean., Univ. of Calif. 6,
27-122.

The Norwegian cyclone models in
relation to heat and cold sources.

Geophys. Publ. Geophys. Norwegica,
2, a =aeoe

On the role of air-sea interaction
in the development of cyclones.
Bull. Amer. Met. Soc., 46, 4-15.

Transformation of polar continental
air to polar maritime air.,
J. Meteor., 2, 94-113.

The warming of Arctic air masses
over the eastern North Atlantic.
Quart. Jour. Roy. Meteor. Soc.,

TT, 355-305.

A vertically integrated wet
diabatic model for the study of
eyclogenesis. Proc. Int'l Symp.
on Num. Wea. Pred., Tokyo, Nov. 7-
13, 1960. 185-204, Met. Soc. of

Japan.

Some results of experiments with
an integrated, wet, diabatic
weather prediction model. Sci.
Rep. No. 2, Contract Nonr-285 (09).
New York University. 28 pages.

On the modification of air-mass
over the Japan Sea when the out-
burst of cold air predominates.
J. Met. Soc. Japan, 35, 311-326.

Winston, J.S.

Martin, D.C.

Bushby, F. H. and M. K.

Malkus, J. S.

Marciano, J. J. and
G. KE. Harbeck

1958:

1959:

1962:

1955:

1958:

1962:

1952:

On the estimation of energy exchange
between the Japan Sea and the
atmosphere during winter based upon
the energy budget of both the
atmosphere and the sea. J. Met. Soc.
Japan, 36, 123-133.

Physical aspects of rapid cyclogenesis
in the Gulf of Alaska. Tellus, 7,
481-500.

The relation between non-adiabatic
heating and the errors of numerical

forecasts. Proc. Int'l. Symp. on
Tokyo, No. 7-13

Num. Wea. Pred.
1960, 253-296. Met. Soc. of Japan.

Further computations of 24-hour
pressure changes based on a two-
parameter model. QJRMS, 81, 396-402.

A graphical prediction model incorpo-
rating a form of non-adiabatic heating.
J. Meteor., 15, 1-8.

Large scale interactions, Ch. 4 in
The Sea, New York, John Wiley and
Sons, pp. 88-94.

Mass transfer studies. U. S. Dept.
of the Interior, Geol. Survey, No.
229 Water Loss Investigations, Vol. I.
Lake Hefner Studies. Tech. Report.

THE THREE-DIMENSIONAL OCEAN CIRCULATION DRIVEN BY DENSITY GRADIENTS

IN AN ENCLOSED BASIN

Kirk Bryan
U. S. Weather Bureau, Washington, D.C.

ABSTRACT

Estimates of poleward transport of heat based on the heat balance of
the ocean surface indicate that ocean currents in the North Atlantic trans-
port from 10-20 percent as much heat poleward as the entire atmosphere in
middle latitudes. Similar measurements for the Pacific based on heat balance
are less reliable. An analysis of hydrographic data obtained during the IGY
and NORPAC expeditions permits the examination of different components of
the heat transport. Of particular interest are the two components associ-
ated with the thermohaline circulation, and the wind-driven subtropical
gyre. The poleward heat transport by these two agencies is of the same
order of magnitude. In the North Atlantic the thermohaline circulation
and the wind-driven gyre both transport heat poleward. On the other hand,
present evidence on the circulation of the North Pacific suggests that
there these two important components tend to cancel each other. The rela-
tive contribution of smaller scale, transient motions is unknown.

A numerical model is proposed to gain further insight into the mech-
anism of poleward heat transport. Solutions are obtained for an enclosed
basin of planetary scale bounded by two parallel meridians. The equations
of the model closely correspond to the complete Navier- Stokes equations
with viscosity and conductivity terms replaced by equivalent terms repre-
senting the effects of small-scale diffusion of momentum and heat, respec-
tively. For the case of no wind, scale analysis suggests that the total
poleward heat transport in the basin should be proportional to

KL2A0*/d

where k is the diffusion coefficient in the vertical,A@*/L is the north-
south temperature gradient imposed at the air-sea interface, and d is the
scale depth of the thermocline. The constant of proportionality obtained
by the numerical calculations is consistent with estimates of poleward
heat transport based on the heat balance method, and empirical determina-
tions of kK.

19
INTRODUCTION

An essential factor in determining the climate of the temperate zone
of the Northern Hemisphere is a strong transfer of heat from the ocean to
the atmosphere during the autumn and winter. Part of this heat (roughly
half in the North Atlantic) has been received by the ocean during the
previous spring and summer. The remainder is supplied by the lateral
transfer of heat by ocean currents from other areas which receive a net
surplus of heat on an annual basis. Detailed studies of the heat balance
of the ocean offer one means of making a quantitative estimate of heat
transfer. In Figure 1 estimates of the poleward transport of heat based
on heat balance maps of Sverdrup (1957), Budyko (1956), and Albrecht (1960)
are compared with direct measurements of energy transport in the atmosphere
made by Starr and White (1954). Many features of the estimates in Figure 1
differ, but there is general agreement that a significant poleward transport
of heat does occur in the Northern Hemisphere, the greater part of which
takes place in the North Atlantic. A discussion of these estimates is
given in an earlier paper (Bryan, 1962).

Most of the present-day knowledge of ocean circulation is based on
detailed measurements of temperature and chemical properties. While this
data is very important in tracing the origin and movement of water masses,
it is difficult to use it directly in studying heat transfer by ocean
currents. Mathematical models are needed to relate ideas gained from
water mass analysis to the heat balance of the ocean and large-scale inter-
action with the atmosphere.

Recently, considerable attention has been devoted to the problem of
the maintenance of the oceanic thermocline (Robinson and Stommel, 1959,
Welander, 1959, Stommel and Webster, 1962, Blandford, 1965). These studies
are directly relevant, since they deal principally with the manner in which
heat is transferred from the surface to lower levels in the ocean. The
steady-state solutions of the thermocline theories are intended to apply
to the subtropical region of ocean basins, away from strong boundary
currents. A disadvantage of these solutions is that they cannot easily
be extended to include subarctic gyres and boundary regions. In particular,
difficulties exist in treating regions in which the stratification is
unstable, or nearly so.

These thermocline investigations form the point of departure for the
present study. Solutions for an entire closed basin are obtained by
numerical methods. To include regions in which convection may ocur, the
vertical heat diffusion coefficient is a constant as in the model of
Robinson and Stommel (1959). For unstable stratification this coefficient
is effectively infinite. Since small, but significant departures from
geostrophy exist in strong currents near lateral boundaries, the model
includes the momentum equations in nearly complete form, without the
geostrophic approximation.

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LY ——— J
JYFHASOWLV

The density anomaly of sea water is due to the distribution of both
salinity and temperature. Source or sink regions at the ocean surface
for salinity coincide with areas in which evaporation exceeds precipitation
or vice versa. With exceptions like the equatorial rain belt, the surface
of tropical oceans is a source region of both heat and salinity. On the
other hand, cooling and an excess of precipitation make the ocean surface
in subarctic areas a sink of both heat and salinity. To simplify the
formulation of the present study of large-scale heat transfer by ocean
currents it will be assumed that the boundary conditions of temperature
and salinity have the same dependence on latitude and are independent of
longitude. Neglecting second order terms the equation of state of the
model is given as,

p = eof 1 - a(T - Te) + o(S - S44

where aand o are the respective expansion coefficients for temperature,
T, and salinity, S. Since T and S obey the same type of conservation law,
and the boundary conditions are proportional, the two variables are no
longer independent. A virtual temperature (Fofonoff, 1962) may be defined
as

@=T- mice
Pow ono )

The single variable, 6, then combines the effect of both T and S on the
density field.

EQUATIONS OF THE MODEL

The basic equations are taken to be the Navier-Stokes equations,
written for a Mercator projection in a rotating frame with the following
assumptions: a) hydrostatic balance, b) variations in density neglected
except where they appear as a coefficient of g (the gravitational accel-—
leration), c) viscosity and conductivity are replaced by simplified terms
representing the diffusion of momentum and heat by smaller scale transient
disturbances.

Let m sec

sin 9

where $is the latitude. If A is the longitude, and a the radius of the
globe, x and y coordinates are defined as follows:

dx = add
dy = amd

A and y are respectively the eddy diffusion coefficients in the horizontal
and vertical.

22
With this notation the equations of motion and continuity are:

u, + muu_ + mvu_ + wue - 2n(Q + uM yy = -m(P/p_) + Ku + Am@V2u (1)
T x y z a ox ZZ

Wy Ss 22 2

vr + muv + INN + wv. + 2n(Q + qu NED + cvee + Am-V<v (2)

9/0 | = (EVD) = (3)

Wiese tS m?[(u/m) + (v/m) 1 (4)

The simplified equation of state used in these computations is then
p = PoC 1 - a0)

The conservation equation for temperature is

8. + mu@ + 6 + 5% 202
t gg YEA YEE, Gian Ars (5)
In (5), 6 is
0 @ <0
6= z
aL 6 >0O

indicating that for stable stratification the vertical mixing is a constant
but for unstable cases effectively infinite.

The boundary conditions are the appropriate ones for a basin bounded
on the east and west by two meridions one radian of longitude apart. To
the north and south the basin is bounded by two parallels of latitude one
radian of latitude apart. The south wall is placed 10° of latitude away
from the equator. Let x, y, z be the three coordinates of an interior
point of the basin. Then

0o<x< xX
o<y< Y

=p < z < 0

The boundary conditions on temperature are such that no heat is diffused
through the lateral walls or the ocean bottom.

6 =0 x = 0,X
x

e, = 0 y = 0,Y

6 =0 z= -D
Zz

Temperature is prescribed at the upper surface as a linear function of
latitude. Let 46* be a scale temperature.

6(¢) = Ab*[ 1 - (6 - o/ ($y = ods z= 0.

The boundary conditions on velocity in this preliminary investigation
filter out external gravity waves, and eliminate any stresses acting at
the bottom or at the upper surface.

WeWU BY 0 mS 0), AWDo

Both the normal and the parallel components of velocity vanish at the
lateral boundaries.

x = 0, X

7 8 O05 i

Equations (1) - (5) are solved by finite differencing using a grid of
19 x 19 points with 6 levels in the vertical. In some cases a more refined
net was used close to the western boundary to obtain a better resolution of
the boundary current. The numerical scheme is based on ideas proposed by
Arakawat/and Lilly (1965). Details of the numerical method will be published
in a separate paper.

RESULTS

In laboratory studies of hydrodynamic models scale analysis is an
essential tool. It is also useful in a numerical study to isolate the
important variables and eliminate redundant calculations. Following ideas
proposed by Robinson (1960), a scale velocity, V*, and a scale depth of the
thermocline may be defined in the following way. In terms of a geostrophic
balance between the vertical variation of velocity and the horizontal
temperature gradient,

2 2 V*/d = g a AO*/L

The requirement that the vertical diffusion of heat be of the same order as
the horizontal advection of heat may be expressed as,

V* Ae*/L = KAe*/d2

1/ Arakawa, "Computational Design for Long Numerical Integrations of the
Equations for Atmospheric Motion," paper presented at the 44th ‘nnual
Meeting, A. G. U., Washington, April 1963.

A definition of V* and d may be obtained by combining these two
relationships.

Scale analysis indicates that there are only three completely independ-
ent variables in the problem. A convenient formulation of these three
dimensionless variables is given below. An estimate of their approximate
magnitude in the case of the real ocean is also indicated.

= V¥D2/(xKL) ~ 100
= V*L/A ~ 10 - 1000
-5
Ro = V*/2QL ~ 10

R, and R, may be considered effective Reynolds numbers for the vertical
and horizontal,respectively. R, is a Rossby number. The estimate of the
magnitude of the Rossby number is for the ocean interior. Much larger
values would be appropriate for the type of flow in the western boundary
current.

A useful nondimensional form of the total poleward heat transport in
the basin is obtained by normalizing the calculated northward flux with
the amount of heat transferred down to greater depths from over an area,
L’, through a vertical temperature gradient of A@*/d.

H/H* _ Poleward Heat Flux

5) KL2 Ae*/d

From general considerations
H/HAm=eeP (Rap) Bay Repent)

@” “©?

Numerical integrations of the model were performed to determine F as a
function of the independent parameters for which the model ocean
settled down to an equilibrium state.

In most cases the initial conditions are a state of uniform strati-
fication and no motion. When a north-south density gradient is imposed
through the surface boundary condition, convection takes place in the
northern part of the basin. This in turn leads to the buildup of horizontal
density gradients in the main body of the fluid. As the parameter, R., is
increased, the effective horizontal mixing decreases. This allows an
increasingly complex flow pattern to form. To resolve these complex patterns
a detailed numerical grid and a large amount of calculation are necessary.
The calculations of this study are therefore restricted to cases in which
Re < 36. Within this range an equilibrium is usually obtained after a
numerical integration over the equivalent of a decade.

e’

The behavior of the heat transport as a function of time is show in
Figure 2 for four different cases. These calculations show the effect of
a four-fold change in the Rossby number with the other parameters kept

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Wath

*g amnesty

~H/HeT

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st ‘p ‘ygdep eTeos ey *s/7ud G = » pue “T °309q. OTXS°%Z = 0
©o8T = x0V Jog ‘(9usT2) olLZ = YX 42 eueTd Teuotptseu

ag} pues ‘(4J59eT) oS€ = > 4e oueTd [Teuoz oy UF sUOTZOES SInqeredmaL, °f sInaTY

SaqNLILV7 AGNLIOSNOT LSVA
OL 0 AGE 7 AOS OY 50S cOv .O0& «Od <OL a0)

H (45°) x 1014 cal/sec

MASS TRANS. (x 10° tons/sec)

Figure 4.

OUIBW fd

Ln K (om SeC —

Above: Poleward heat transport estimated from the formula
H = .2H*, as a function of A6* and K . Below: The
associated strength of the thermohaline circulation.

Circuit Time (Centuries)

constant. For a value of x equal to 5 em@/s the change in Rossby number
would correspond to a change in the north-south temperature contrast from
92 +o 36°. Note that this large change does not appear to have a corres-
pondingly large effect on the nondimensional heat transport. The three
solid curves in Figure 2 are for the cases in which R, is equal to 1000.
The dotted curve represents a single test calculation made for Ro equal to
2x10-5 and Rj equal to 100. The total depth, D, appears only in R,. The
test calculation shows that beyond a certain point, the purely thermal
solution is insensitive to the total depth. A similar result has been
obtained previously in the thermocline calculations of Stommel and Webster

(1962).

Figure 3 shows vertical sections made for a zonal and meridional plane

cutting the basin. The temperature has been normalized by dividing it by

Aex . Note that the isotherms are fairly flat over most of the basin.
Exceptions occur near the western boundary in a narrow zone, and in the
northern part of the basin. The upturned isotherms near the western wall
are associated with an intense, northward moving boundary current. A much
slower, but deeper compensating current moving southward exists below. This
western boundary current differs from that of the wind-driven ocean theories
(Stommel, 1948) in that the net, vertically integrated mass transport is
zero. This type of boundary current associated with the thermohaline circula-
tion has been anticipated by Stommel (1958, page 157) in his prediction of
an undercurrent in the vicinity of the Gulf Stream. Analytic solutions have
been obtained only from a simplified linear model by Takano (1962).

Another set of calculations similar to those shown in Figure 2 indicate
that H/H* depends markedly on the Reynolds number only in the range
O<R, < 10. For larger Reynolds numbers horizontal mixing plays a rather
small role in the poleward heat flux. Through an extrapolation of the
results, it is estimated that the equilibrium value of H/H* for very large
values of Reynolds numbers would be approximately .2, assuming that R, = 100
and the Rossby number is in the geophysical range. The oceanographic
interpretation of this result is shown in Figure 4. Assuming that H/H* is
0.2, the total poleward heat flux is given as a function of the vertical
diffusion coefficient, «x , and the total meridional temperature difference
imposed at the surface. The Atlantic Ocean is known to have a direct
thermal-haline circulation. For a rough comparison of heat transport in
the model with observations we note from Figure 1 that the poleward flux
at 45° latitude in the North Atlantic is about 2x10" cal/s. Allowing for
the effect of salinity, a north-south virtual temperature difference of
18°C is in best agreement with surfacg temperature data. From Figure 4 we
see that a vertical diffusion of ance /s would be required for the model to
have a poleward heat flux of 2x10°" cal/s. This is a reasonable value of
k , Since independent empirical estimates based on water mass analysis are
all of the order of unity (Robinson and Stommel, 1959).

In the lower part of Figure 4 the strength of the thermohaline circula-
tion is plotted, also based on an extrapolation of the numerical results to
the case of very high Reynolds numbers. The total rate of overturning in

a vertical meridional plane for a temperature difference of 18°¢ ana «

equal to 5 em@/s is 40 million tons/s. This is about 1/2 the observed
transport of the Gulf Stream (Stommel, 1958). The average circuit time for
water to sink to great depth and rise to the surface again is obtained by
dividing the strength of the circulation into the volume of a basin 5 km
deep. The right hand ordinate of Figure 4 indicates that the circuit time
is of the order of centuries for the particular case under discussion.

Further calculations are in progress to test the effect of much larger
variations in the Rossby number, and the modifications introduced by the
effect of wind acting at the surface. It is hoped that such calculations
will bridge the gap between theories of the thermocline and theories of
purely wind-driven circulations. In principle there are no difficulties in
including salinity in the model which will allow a much more realistic
formulation of the boundary conditions. Robinson and Stommel (1959) have
emphasized the importance of the vertical diffusion, «, and have pointed
out how little is known about the forced convection represented by this
parameter. Based on the simplified density-driven model of this study, «
is shown to be the principal factor in determining the partitioning of
poleward heat flux between the hydrosphere and the atmosphere.

REFERENCES
Albrecht, F. 1960 Ber. Deut. Wetterdienstes, 66,
Bd. 9.
Blandford, R. 1965 J. Mar. Res. (In Press).
Bryan, K. 1962 J. Geophys. Res., 67, p- 3403.
Budyko, M. I. 1956 Atlas Teplovoga Balansa, Moscow
Fofonoff, N. 1962 The Sea, Vol I, M. N. Hill, Eda.,
Wiley, N. Y., London, p. 368
mE, Wo Kee 1965 Mon. Weather Rev., 93, p- 11
Robinson, A. R. and Stommel H. 1959 Tellus, 11, p. 295
Robinson, A. R. 1960 Deep-Sea Res., 6, p. 311
Starr, V. P. and White, R. M. 1954 Geophys. Res. Dir. Paper, 35,
Bedford, Mass., U. S. A.
Stommel, H. 1948 Trans. Amer. Geophys. Un., 29,
p. 202
1958 The Gulf Stream, Cambridge Univ.
Press
Stommel, H. and Webster, J. 1962 J. Mar. Res., 20, p. 42
Sverdrup, H. U. 1947 Proc. Nat. Acad. Sci., 33,
jo Sls)
1957 Handbuch der Physik, 48,
Springer-Verlag, Berlin
Takano, K. 1962 Records Ocean. Works Japan, 6,
De 60
Welander, P. 1959 Tellus, ll, p. 309

ON THE PRESENT STATE OF KNOWLEDGE IN ATR-SEA BOUNDARY LAYER PROBLEMS

H. U. Roll
Florida State University
Department of Meteorology

INTRODUCTION

I hope that the title chosen for this talk has already indicated with
sufficient clearness that I am going to deal with processes of small scale.
We all are very well aware of the fact that air-sea interaction is not
restricted to such small-scale processes but extends through the whole scale
of motions comprising mesoscale and synoptic processes and reaching even the
planetary scale by affecting the atmospheric circulation and the energy
balance of the earth. Nevertheless, the limitation imposed on this lecture
helps to focus our attention on the crucial region of air-sea interchange.
This comparatively shallow layer with a thickness of only a few meters in
air and water and characterized by vertical fluxes and energy transforma-
tions of different kinds apparently holds a key position in the interaction
between the atmosphere and the ocean. All the motions and processes of
other scales and related to air-sea interaction are in some way predetermined
by the small-scale exchange occurring in the boundary layer air-sea.
Therefore, any progress in our understanding of the interaction between ocean
and atmosphere on the whole and in all its different parts cannot be
accomplished without a simultaneous or preceding progress of our knowledge
about the physics of this interchange in this shallow boundary layer.

This can be stated much more easily than it can be translated into
action. The sea surface is distinguished from the atmospheric boundary
conditions prevailing over land by very peculiar properties. On the con -
tinents, shape and size of the elements of surface roughness are clearly
defined and comparatively easy to determine. Generally, their nature and
locality are fixed and they neither vary with time nor do they depend
strongly on atmospheric conditions. Their aerodynamics are relatively well
defined and known.

Contrary to the boundary conditions found over land the surface rough-
ness encountered at sea is composed of a great variety of moving elevations
which are different in size, shape, and velocity as well as subject to
continuous and irregular changes. The dimensions, the spatial distribution,
and the temporal variations of the ocean waves are governed by statistical
laws wherein the character and speed of the air flow play a decisive role.
Moreover, the wind generates orbital motion and drift current in the sea
and it is quite obvious that these water movements will react on the air
flow. With increasing wind speed, the formation of foam and spray, which
implies a disintegration of the sea surface, affects large areas and extends
to a certain height, thus creating a transition zone between air and sea.
Therefore we must realize that the boundary region between air and sea is
an extremely variable, ill-defined, and hardly accessible zone where the
coupling between atmosphere and ocean occurs in a very complicated manner.

These dynamic properties of the sea surface considerably increase the
difficulties inherent in any investigation concerned with the mechanism of
the air-sea boundary layer.

With a view to this severe handicap it is only of little comfort that,
on the other hand, the sea surface also has a few pleasant properties. The
local differences, which are most prominent over land, are substantially
reduced on the oceans, these being much more uniform in this respect than
the continents are. Their capability of acting as sources or sinks for
heat and moisture shows but little variation from one place to another.

For example, it has been recently reported by Brocks (1963) that, in the
southern part of the North Sea, the correlation between simultaneous
measurements of air temperature as well as between that of humidity or wind
speed executed within a sea area of at least 25 nautical miles was found

to range mostly between 0.9 and 1.0, which proves the high degree of spatial
homogeneity at least for this area. Further, since the diurnal and annual
variations are much smaller than those on land, there is also a pronounced
uniformity in time at sea. Thus, in some respect, the oceans offer an

ideal field for meteorological investigation provided that it is possible

to overcome the experimental and theoretical difficulties mentioned before.

Every review must have a certain reference level from where it starts,
i.e. a certain amount of knowledge which can be taken for granted, since
it is impossible to give a complete treatment within a rather short time.
Such a reference level can best be provided by a suitable publication.
I am in the happy position of being able to make reference to two excellent
reviews on our present subject. The first, given by E. L. Deacon and E. K.
Webb (1962) provides a very concise and still detailed treatment of small-
scale interactions air-sea. The viewpoint of the second review, which has
been elaborated by 8 distinguished scientists (Benton et al., 1962) is more
general, its main object obviously being to put the finger in the wound of
insufficient knowledge and to show what should be done about it.

I shall take these two publications as a base for my discussion assuming
that the state of affairs as it is reported therein is more or less known
to the audience.

The interchange occurring in the boundary layer air-sea is manifold.
When attempting to treat it systematically we may perhaps make a subdivi-
sion by separating

the exchange of energy from

the exchange of matter and from

the exchange of electrical charge.
Although the transfer of matter and of electrical charge through the
boundary layer certainly has interesting or even fascinating aspects, I
would rather like to confine my discussion to the exchange of energy thereby

including the exchange of water which, owing to the latent heat of vaporiza-
tion, must be considered as a - quite important - part of the energy transfer.

Primarily,there are four ways in which energy can be exchanged between
the oceans and the atmosphere:

(1) by the transfer of momentum,
(2) by the radiative interaction,
(3) by conduction and convection of sensible heat, and

(4) by molecular and turbulent transport of latent heat in the
form of water vapor.

I would like to deal with these different kinds of energy exchange in
the order indicated above. When doing so I should emphasize that a complete
review cannot be expected, because here in Tallahassee I do not have at
hand my extensive and detailed file of references which, of course, I could
not bring with me. Therefore, my presentation is certainly biased by a fair
amount of randomness as far as the literature reported is concerned.

THE TRANSFER OF MOMENTUM

With a view to the well-known irregularity of motion near and at the
sea surface our final aim must be to get a complete time record of the
field of motion both in air and in water as well as of the distribution of
pressure and stress in the marine boundary layer. These time records must
show as high an amount of temporal resolution and must also be as long as
would be necessary in order to allow (1) a spectral analysis of all fluctua-
tions which may contribute to the transfer of momentum and (2) a reliable
estimate of the vertical momentum flux. Further, a time record taken at
only one point would certainly not be sufficient but must be supplemented
by others taken nearby or by some suitable device which provides informa-
tion about the spatial properties of the flow. On the base of such empirical
information and using sound physical principles, theory must try to develop
suitable models which can be applied for predicting purposes.

Looking first at some empirical evidence on the instantaneous wind
field around moving ocean waves we can hardly see anything at all. The
only measurements, which came to my knowledge and which at least supply a
certain part of the information wanted, are those published by Pond, Stewart,
and Burling (1963) and the - still unpublished - results obtained by Brocks
and Hasse (1963).

Pond, Stewart, and Burling measured turbulence spectra of the
"downstream" component of the wind over waves of approximately 30 cm height
using a hot-wire anemometer. The probe was mounted at 1 to 2 m above the
water level, at which height the mean wind speed was about 3 m/sec. No
indication is given as to whether the measuring site was close to the shore
or well on the open sea. But the smallness of the wave heights mentioned

therein leads us to believe that the anemometer was mounted on a fixed
construction near the shore. The result obtained was presented in the form
of a one-dimensional energy spectrum (Figure 1) giving the energy of the

fog a(k) (cgs)

0 0180-0210
+ 0190-0150
+ 0150-0200
+ 0200-0210

fog hk (cgs)

Figure 1. Energy spectra of the downstream component of the wind velocity
fluctuation at a height of 1 to 2 m above the sea surface as a function
of the wave number k = 2nf/u. Log-log plot. The run of 30 min.
duration is broken down into subsections of 10 min. each to show the
steadiness of the spectra. The straight line has a slope of -5/3 (from
Pond, Stewart, and Burling, 1963).

fluctuations in the dowmstream wind component as a function of wave number k
where

IS SB Ay s/w . (f =frequency of fluctuations, u= mean wind speed)

The straight line corresponds to Kolmogoroff's theory of local isotropy
and - in the double-logarithmic graph - has a slope of -5/3 which says that
spectral energy density function goes with the -5/3 power of the wave
number. As it can be taken from the graph, the results provide further
support for Kolmogoroff's contention that there exists a universal form

to the high number part of the spectrum of high Reynolds number turbulence.

A similar - as yet unpublished - result has kindly been communicated
to me by Brock and Hasse (1963) who recorded the horizontal and vertical
components of the wind speed and the air temperature as well. These
measurements were made by means of a buoy (Figure 2) carrying a stabilized
mast on which hot-wire anemometers and platinum resistance thermometers
as well as vertical accelerometer were mounted. The measuring site was
well away from land and - owing to the distance between buoy and research

Figure 2. Buoy with gyro-stabilized mast carrying sensors for recording
the fluctuations in horizontal and vertical wind components and in air
temperature. In the foreground:small buoy for recording the wind speed
close to the sea surface. In the background: research vessel "Hermann
Wattenberg" of the Oceanographic Institute at Kiel University connected
by floating cable with the buoy. (By courtesy of Dr. K. Brocks.)

ship being about 250 m = also any disturbing influence from the ship was
avoided. Brocks and Hasse computed variance spectra for the horizontal

and vertical wind components and for the air temperature which again seem

to support Kolmogoroff's -5/3 power relationship, apart from some deviations
in the wind fluctuations at frequencies of 0.3 to 0.4 c/sec, which certainly
originate from sea waves. Their results are reproduced in Figure 3 where
the products of spectral intensity and frequency are plotted as functions

of the spectral frequency f for the horizontal and vertical wind components
u, w, for the air temperature @ as well as for the covariances u'w' and

6'w'.

Figure 3.

Q05 qv Q2 0.5 1 2 5 10 20 ¢(Hz]

FE (1) [m?sec’]

VARIANZSPEK TRUM
14 OSTSEE 1962
1.2 21062 12.50Uhr 120 sec

472095 G0: 5.7 misec
1.0 Pian OO
ets ee ter Oe ew) a es Trend etiminiert
06
Qe

HORIZON TALWIND

a2
00 =

1-E(t) [msec 4]

0.12
PRODUKT u'w

VERTIKALWIND
Qo

PRODUKT ®6'w'

f-E(t) [°c]?

Ay TEMPERATUR

1-10-

ty)
005) 02 as 1 2 5 10-20 _ ¢ (Hz

Variance spectra of horizontal wind component u,

covariance u'w'

vertical wind component w,
covariance 6'w'

air temperature 6.

The products of spectral intensities and frequency f are plotted versus
spectral frequency f (Hz =c/sec). (By courtesy of Dr. K. Brocks.)

So far this is all that has come to my notice about measurements of
the instantaneous wind field around ocean waves. The spectral analysis
seems to be a suitable method of representing and studying such very ir-
regular motions. However, I would wish that it may not only be applied
to the wind field over the sea but also to the wave motion at and - if
possible - below the sea surface. Simultaneously taken records of this
sort would yield highly useful information on the mechanical interaction
between air and sea, in particular if the measurements of the motion in
both media were supplemented by records of the pressure distribution and
its fluctuations. Apparently, plans and preparations for such an approach
as well as preliminary field tests are being made at the University of
British Columbia, Vancouver (Stewart and Burling, 1961).

Certainly the whole problem is easier to tackle by laboratory
meastrements, although the result obtained there may not always be meaningful
with respect to open-sea conditions. One interesting paper of this kind
has recently been published by Schooley (1963) who tried to measure the
wind field above wind-generated water waves in a short tunnel by photograph-
ing the tracks of neutrally buoyant soap bubbles. The data could be sum-
marized in form of vertical wind profiles (Figure 4) above certain points

3)
F¢
ied
=
w
ec
WwW
q
=
WwW
>
3
a
c<¢
uJ
oO
z
<
w
a

5
OS aOMNIIIIZ=ON sero UnESINK4 EES ING NAATANNO NNO NNO II2=ONImN2
DISTANCE ALONG WAVE (em)

WAVE HEIGHT (cm)

Figure 4. Vertical wind profiles above four different points along a
water wave in a wind-water tunnel (from Schooley, 1963).

along the wave profile and show the expected strong increase of wind speed

with height above the crest region as well as the less steep gradient above

the trough. I said "expected" because a similar result had be#A% obtained “hree
decade@ ago by Motzfeld (1937) who investigated the air flow over a wavy

te)

but solid wall in the institute of Prandtl. According to Schooley's find-
ings (Figure 5) the flow has a maximum speed which occurs at about 1.5 to
2.5 em above the water surface. This ‘jet" effect, as it is called by
Schooley, points to a systematic deviation from the log-profile at a level
of about 1.5 to 2.5 wave heights above the surface. This could be of
importance as will be explained later.

DISTANCE ABOVE WATER SURFACE (cm)

Gees
Peete
aoe
hn il pee See
0 boa
Ol i Oo.) Si hae aS Gum nO aS) Omen tle

HORIZONTAL WIND SPEED (METERS /sec)

Figure 5. Vertical wind profiles above four different points along a
water wave showing a maximum velocity just above the boundary layer.
(from Schooley, 1963)

On the whole, laboratory measurements seem to be very promising, in
particular for developing and checking theoretical models. I have been
told that a study on the instantaneous wind field around moving water waves
is being mede at the National Center for Atmospheric Research by means of

@ rather sophisticated equipment. The results aspired to will be of great
interest.

As long as measurements of the instantaneous wind field over the ocean
waves are not available, the average vertical wind profile should at least
furnish us some useful information which can be interpreted in the light
of the turbulence theory developed and checked by means of measurements in
laboratories or over land. Now a real trouble begins! I would not like to
spend much time describing the observational problems, but let me only
mention this: There are two possibilities of fixing the height of a certain
mean wind speed:

(1) One can take the average distance from the wavy sea surface.
This can be done by placing the anemometer on a fixed construction or on a
Ploating base that does not participate substantially in the wave motion.
No measurements in the wave troughs are taken in this case.

(2) The measurements are made at a point that has the same distance
from the wave sea surface at any instant, i.e. the anemometer oscillates
with the sea surface. This can be realized by using a float or a buoy as
carrier of the instrument. In this case, measurements may include the
trough region.

‘Up to now, it is not clear which procedure gives the better estimate of the

mean Wind profile. Naturally, the influences coming from the fixing of the

zero level will only be significant in the immediate vicinity of the waves.

Unfortunately, this is the very height range where the vertical wind profile
is of particular interest and importance.

Apart from such observational difficulties, there are the problems of
interpretation of the results. During the last years we were very happy
that the majority of the vertical wind profiles measured showed a log-
distribution which can be easily interpreted in terms of the turbulent
boundary layer theory. Out of 26 studies I reviewed, 1} reported a log-
profile, 3 could explain their deviations from the log-profile by the
influence of thermal stratification, 3 did not say anything about it,
because the wind speed was only measured at 2 levels (which certainly is
the easiest way) and only 6 papers, mostly published before 1950, reported
@& pronounced deviation from the log-profile in the lowest levels (below
2m). The occurrence of such a "kink" in the wind profile could, however,
be explained by observational errors (determination of heights) or by dis-
turbances originating from the carrier of the instruments, e.g. from the
ship, float, etc. Thus, the validity of the log-profile appeared to be
well established also for the marine boundary layer under adiabatic con-
ditions, and many conclusions about the aerodynamic roughness and the
friction coefficient of the sea surface as well as the wind stress at the
sea surface were based on this fact.

Quite recently, however, some doubt has been shed on the validity of
the log-profile for representing the mean vertical wind distribution over
the sea. For instance, Takeda (1963) and others found a kink in the
lower part of every log-profile they measured over the sea and, after having

checked every possible influence carefully, stated that this deviation was
not caused by any instrumental or observational error. They were led to
believe that the structure of the air flow over the undulating sea surface
is different from that over land or along a solid wall. Theoretical argu-
ments (Miles, 1957; Stewart, 1961) also suggest the existence of a critical
layer in the wind profile over waves where the wind speed is equal to the
phase velocity of the waves. Stewart predicts a nonturbulent, organized
and wave-like motion below that level which is connected with a reduction
of the turbulent stress and wind shear. At present it is yet too early to
interpret the kink recently found in vertical wind profiles over the sea by
referring to Miles' and Stewart's critical level. A systematic investiga-
tion of the fluctuations of flow, both immediately above the sea surface
and at it, is necessary in order to bring this problem nearer to solution.

Before leaving this subject of vertical wind profiles over water let
us cast a short glance at a diagram ( Figure 6) which summarizes the results
obtained from log-profiles by applying the turbulent boundary layer concept.
Under adiabatic conditions those profiles yield corresponding values for
the aerodynamic roughness z,) and for the friction velocity ux which is
defined as the square root of the ratio surface wind stress aby air
density p.

In the diagram z) is plotted as a function of ux. We are confronted
with a very confusing result, because some evidence for a decrease of
with growing u, can-be found as well as some proof for its increase ae
growing ux or its constancy. Thus we must state that this relationship
is by no means well understood at present. Even the physical meaning of
the so-called roughness parameter Zq is obscure. In the boundary layer
theory Z describes the scale of turbulence at the level where the mean
wind speed 7 is equal to zero. Remember the well-known log-profile of

wind speed a ve

=i In (e)

u Daa aan (1)
where 1 =O for z =0 and the turbulence present at this level is described
by the mixing length 1 =k z, (k = von Karman constant). At sea there is,
in general, no level at Gracin the mean wind velocity U =0O, since the water
surface itself may move with appreciable speed (by about 4 percent of the
wind speed taken at 10 m). Thus, the boundary layer model needs considerable
amendment and refinement in order to be applicable to the complicated
mechanism of air-sea interaction.

The results presented so far referred to the adiabatic wind profile.
Regarding the wind profile under nonadiabatic stratification very little
evidence is available from the sea which can be compared with the theoretical
approaches given by Monin and Obukhov (1954), Ellison (1957), Yamamoto (1959),
and Panofsky, Blackadar and McVehil (1960). The reason for this is that,
in order to be able to apply these theories, data on the vertical heat flux
are needed apart from the simultaneous measurement of the vertical momentum
flux and wind profile. It is very difficult to get reliable information

DYNAMIC ROUGHNESS Zo

CM
1O-!

10-2

107”

BROCKS ('59) a
(BALTIC)

PORTMAN (1960)

0) 20 30 8640 50
FRICTION VELOCITY Uy CM/SEC

e sea surface as a
Summarizing graph.

9 of th

ty uy-

Dynamic roughness z

function of friction veloci

Figure 6.

about these quantities over the sea. The small amount of information
available seems to indicate that at sea the influence of thermal stratifica-
tion on the wind profile can be taken into account in the same manner as
this is done over land (Deacon, 1962).

In this connection the following seems to be important: The stability
of the air above the sea does not only depend on the vertical temperature
distribution. It is determined by the vertical variation of density, i.e.
there may be an effect of the humidity gradient also. Kraus (1964) has
drawn our attention to the fact that a strong humidity decrease with height
in the lowest layer above the sea may even be sufficient to.reverse the
stabilizing effect of a small temperature increase with height. A decrease
of water vapor pressure of 5 mb would compensate a temperature increase of
0.5°C. Therefore, according to Kraus, this effect must be taken into
account under relevant circumstances, for example, by an additional term
in the Richardson number which is normally used as a measure of stability.

So far I have been talking about the wind field in the marine boundary
layer. The quantity that is of most importance in the field of mechanical
interaction air-sea certainly is the wind stress acting on the sea surface.
Very little is know as yet about the normal wind stress, its size and
spectral distribution associated with the turbulent wind blowing over the
water. This lack of knowledge is regrettable, as it seems we may be sure
that the atmospheric pressure fluctuations play an important part in the
generation of wind waves at the sea surface. Thus, we have not been able
to check the theory advanced by Phillips (1957, 1958, 1962) who considered
the initial waves generated by a resonance mechanism between the surface
wave modes and the random pressure fluctuation associated with the turbulent
wind blowing over the water and convected by the mean flow. There is, how- -
ever, some very recent empirical evidence (Snyder, 1965) by which the
importance of the resonance mechanism is questioned.

{

The tangential wind stress, however, has been the subject of consider-
able number of investigations, although one cannot say that all the problems
connected with it have been solved. The tangential wind stress is equivalent
to the vertical transport of horizontal momentum in a viscid fluid. In the
turbulent boundary layer of the atmosphere the shear stress is usually
considered as constant with height. Over the sea this may not be true as
we know from the theories of Miles and Stewart, but up to now the vertical
constancy of the wind stress is the generally accepted practice also for the
marine boundary layer. Consequently, the tangential wind stress T observed
in the boundary layer is equal to the tangential wind stress T= T
exerted by the wind on the sea surface. The latter quantity is of fnportance
for quite a number of air-sea interaction problems, e.g. generation and
growth of ocean waves, of drift currents, and storm surges.

It is customary to express the surface drag T, of the wind at the sea
surface in terms of the mean speed Uj at the height 10 m

a)
t =
Che ee 10 M10 (2)

the factor of proportionality Cio being a dimensionless height-dependent
quantity, termed resistance, drag, shear-stress, or friction coefficient.
The problem of determining the surface stress ks is then reduced to ascer-
taining reliable values of Cio:

Estimates of the drag coefficient have been based largely on indirect
evidence. The following five methods have been used so far:

In air: (1) Wind profile method: Under adiabatic conditions the drag co-
efficient can be easily calculated from the log profile. In fact, the drag
coefficient is a function of the roughness length zp and also of the height z.
With a diabatic wind profile, the additional knowledge of the vertical heat
flux is necessary. So far the wind profile method was used in about 22
studies in the field and in the laboratory and supplied quite useful results.

(2) Geostrophic departure method: The covariance - pu'w' of
the turbulent fluctuations in the horizontal and vertical wind components is
“recorded and supplies an estimate for the turbulent Reynolds stress. This is
a rather direct approach. Unfortunately, it needs a fixed or stabilized plat-
form as well as sensing elements of sufficiently rapid response. Therefore,
we have as yet only four or five studies of this kind. The data show a

considerable scatter.

In water:(3) Sea surface tilt method: If an enclosed body of water is
available and the wind has blown for a sufficiently long period as to assure
steady state conditions, then the surface wind stress is assumed just to
balance the hydrostatic forces due to the tilt of the surface. The surface
slope provides an estimate for the wind stress or the drag coefficient.
Quite a number of studies (18) were made up to now, partly in the field,
partly in the laboratory. The necessary accuracy (107') could mostly not

be achieved with small wind speeds. Disturbing effects as stratification in
the water, horizontal density gradients, near-shore effects due to waves,
nonsteady state etc., may make the result uncertain.

. At the water

surface: (4) Surface film method: An insoluble monolayer is applied to
the water surface. Its contraction under the action of the wind provides
a measure of the wind stress. This seems to be essentially a laboratory
method. Only one paper (Vines, 1959) has become know so far. E440

I would like to present the results obtained in a somewhat condensed Oy
form by showing a diagram (Figure 7) containing all the empirical relation-
ships suggested between C,, and U,,. There is, of course, a substantial
scatter in the single measurements which are not reproduced here. Looking
at these different results we are in a similar position as we were before
with regard to the roughness length. There is no satisfactory agreement

.005

Cio

.004

alts T

RELATIONSHIPS SUGGESTED
FOR THE DRAG COEFFICIENT Cio
AS A FUNCTION OF
WIND SPEED Ujo

7 ¢ ets
BROCKS (1959) ZZ pe ROOM nce NEN
(BALTIC) Zi ae Pas c
009° ee St —— | ee
Fi spre Fk (1955) (VAN DORN)
i ee BROCKS (1959) a
Za (NORTH SEA)

Uio

ES eae | es ane
6 8 10 l2 9 16 18

Figure 7. Relationship suggested for the drag coefficient Cj

of the sea surface, as a function of the wind speed
(at 10 m) ujo, in m/sec.

between the different authors and methods. The strong increase of the
friction coefficient with decreasing wind speed, which characterizes
Neumann's result, is now generally assumed to be biased by data of insuf-
ficent accuracy, however. Then the problem remains whether increase with
growing speed or constancy is correct. During the last years, a certain
tendency could be observed to diminish the slope and so to approach the
eontancy proposed by Brocks.

These values refer (or should refer) to adiabatic conditions. The
question arises whether there is an influence of thermal stratification
on the wind stress. Visual observations are in favor of such an effect.
The ocean waves appear to be higher and steeper, the production of foam
and spray is more intense with cold air over warm watér than vice versa. {
Some relevant evidence was reported by Darbyshire_(1955) who obtained
stress coefficients that, for a given wind speed, were twice as great in Dart
unstable cases than in stable ones. More convincing measurements were
reported by Garstang (1965). This stability effect can also be calculated
with the help of the theories of Monin and Obukhov and of Ellison. We
then may express the drag coefficient C, at the level z =a by the relation

k2
Ca a at z (3)
{In——* + a Rif2

Z
°

where k is von Karman constant,a = constant ( = 3.7), Ri = Richardson num-
ber.
This relationship remains to be checked by suitable measurements.

Regarding the possible influence of the fetch on the wind stress,
there is no uniform result up to now. A few scientists found favorable
evidence, whereas the majority could not verify such an effect. There is
some reason to believe that the wind set-up measurements which indicated
such an influence of fetch were biased by coastal wave effects. =

- The resuits reported so far on the variation of the drag coefficient
with wind speed are empirical. There is only one theoretical approach,
namely Munk's (1955) interpretation of Van Dorn's (1953) wind set-up data
which is based on Jeffreys' "sheltering hypothesis." Herewith the existence
_ of sheltered regions with eddies to the leeward of the wave crests is as-
sumed, which implies a phase lag between the wave profile and the pressure
distribution. Under the assurption that the Neumann spectrum is valid for
the wave energy density Munk (1955) computed the form drag caused by a
fully developed sea and found that the young high-frequency waves contribute
much more to wave slope and form drag than the low-frequency waves do, which
mainly determine the elevation statistics. Consequently, these low-frequency
waves seem to be of little significance for the aerodynamics of the sea
surface, a conclusion which confirms the findings of other scientists who
were led to believe that the form drag is principally caused by the small,
slowly moving ripples and wavelets. This would also explain that the drag
is much less affected by limitations in fetch and duration than the wave

height, since the high-frequency waves which determine the form drag reach
their final state much more rapidly than the low-frequency waves dominating
the wave amplitude.

For the composite wind stress, which is composed of the tangential
stress and the form drag, Munk obtained a cubic relationship with the wind
speed. This would imply that the total drag coefficient increases linearly
with wind speed. (The corresponding straight line is entered in Figure Te)
Thus, at least some theoretical support is provided for this kind of
variation of the drag coefficient with wind speed.

On the whole, however, the state reached is far from being satisfactory.
There is a distinct gap between empirical and theoretical work. On one side,
theory is not yet able to interpret empirical findings sufficiently. On
the other hand, theoreticians penetrate with their reasoning into regions
which are not yet accessible to measuring. In the latter respect, I am
thinking of quite a number of important papers concerned with the problem of
wave generation (e.g. those published by Phillips (1957, 1958, 1962), Miles
(1957, 1959, 1962) Hasselmann (1960), and others) which cannot be discussed
here. The same is true with regard to the theoretical work recently done
by Schmitz (1962) who studied the transfer of mechanical energy at the sea
surface. He found for instance, that the equality of the vectors of mean
flow on both sides of the sea surface is not a necessary condition (but, of
course, a sufficient one) for the continuous transfer of mechanical energy.
as it is usually thought. This can also be fulfilled if the mean values of
flow in air and water immediately at the surface differ from each other,

i.e. if the air glides over the water.

Thus, a concentrated effort, made both by experimenters and theoreticians,
is needed in order to clarify the complicated process of mechenical inter-
action at the sea surface.

THE RADIATIVE INTERACTION

Passing now to a short treatment of the radiative processes in the
marine boundary layer we enter a region which can perhaps be characterized
best by the statement that here the importance of the subject is only sur-
passed by the imperfection and incompleteness of our knowledge about it.
Although radiation forms the primary energy source of all processes in
atmosphere and ocean, there is an almost complete lack of data measured in
the boundary layer air-sea which however are indispensable when the radiative
interaction between both media shall be studied. First we need a sufficient
coverage of the oceans with climatological radiation data, but in addition
to that some current information on radiation for the study of synoptic-
scale processes will be necessary, because the empirical formula describing
e.g. the effect of cloudiness on the different components of the radiative
budget deviate substantially from each other. Laevastu (1960), has certainly
taken great care in checking the diverse empirical relationships but finally
it results that he gets a higher net radiation for the cloudiness of 4/10
than for a cloudless sky. This is caused by the fact that he assumes a

linear decrease of the effective back radiation with cloudiness while this
decrease goes with the cube in the case of the incoming short-wave radiation.
This example throws some light on the most difficult problem we are faced
with: to develop formulae for the different components of the radiation
balance in which the influence of cloudiness is adequately assessed. There
may be some doubt about whether this will ever be possible with the present
system of estimating only the amount and kind of clouds and the altitude of
their base. In any case, much more routine measurements of the different
components of radiation are urgently needed at sea.

Since more than 15 years ago a remarkable network of ocean weather ships
has been working continuously and at fixed positions in the North Atlantic
and the North Pacific Oceans. These stations have assembled an amount of
meteorological data which are unique in marine metecrology. Unfortunately,
radiation measurements have been included in the observational program of
some of those stations but during the last years, practically since the
International Geophysical Year (IGY). Now some relevant publications have
come up. I would like to mention that of Ashburn (1963) which presents
daily and mean monthly data on total incoming radiation from sun and sky for
2 years at the North Pacific ocean station "Papa" as well as some values of
net radiation. With regard to the influence of cloudiness a new formula
(the llth as far as I can see) is offered which seems to represent the
measured data better than the other ones. This is but a beginning. Many
additional measurements will be necessary to determine the precision of the
measurements and to develop an equation expressing the total incoming radia-
tion received at the ocean surface as a function of the known physical
variables such as cloud types, cloud thickness, solar zenith distance and
albedo of the sea surface.

A more sophisticated approach regarding the influence of cloudiness
was made by Lumb (1964). Using radiation measurements made on British
ocean weather ships he suggested a relationship for the total incoming
radiation during short periods (1 hour) depending on different cloud
categories and the average solar altitude during that period.

These first results should be warmly welcomed but much more radiation
measuring must be done on the oceans - both close to the sea surface and
also at upper levels = before we will have sufficient evidence about the
contribution of radiative fluxes to the energy balance of the atmosphere
and the ocean surface and before we will arrive at a marine climatology of
radiation which is based on really measured values. Certainly, the
instrumental and operational difficulties are great but not insurmountable
if scientists and engineers would recognize the challenge of this task.
Substantial progress could then be achieved. In addition to radiometers
installed on ships and buoys, such instruments could be carried by planes
and satellites and thus fill the gaps between the necessarily wide network
of radiation measuring stations in the marine boundary layer. Relevant
investigations have already been made. I would like to refer to Clarke
(1963) who presented observations of the long-wave radiative flux in the

50
BSRS ESTs TIL Bea Tasison Ur wedietion data taken by the PROS “7

satellites.
SENSIBLE AND LATENT HEAT EXCHANGE

Considering the exchange of sensible and latent heat at the sea surface
I think the present situation here is somewhat better than it is in the field
of radiation, although it is not entirely satisfactory. The chief aim being
to furnish reliable values of the turbulent fluxes of sensible and latent heat
between ocean and atmosphere, we must try to derive formulae by means of which
these transfer quantities can be calculated with sufficient reliability. It
would be particularly welcome if such computations could be done by only using
the meteorological elements which are usually measured on a routine basis at
sea.

Remarkable success has already been achieved with relatively simple
procedures. I am thinking of the so-called "bulk aerodynamice formulae" for
the vertical fluxes of sensible and latent heat through the marine boundary
layer which have been widely used, among others by Jacobs (1951) for his
road-showing study on the climatology of air-sea interaction. However,
these formulae are not without any problems. In addition, there are funda-
mental questions as e.g. the role of the molecular transport at the sea
surface, the effect of sea spray, the influence of surface films, which
require detailed consideration.

Let me first give a discussion of the bulk aerodynamic formulae. They

are mostly derived in such a way that the ratios of the fluxes H (heat),
E(evaporation), and t (momentum) are formed

H/T = ~¢,(38/3z)/( du/dz) and E/t = -(dq/dz) - (au/az) io

(c_ = spec. heat at constant pressure) which can be obtained by dividing
the equations of definition of these fluxes and assuming that the turbulent
exchange coefficients for momentum, heat, and moisture are equal. This is
already a questionable assumption. Immediately at the sea surface, where
molecular transfer comes into play, these exchange coefficients are certainly
not equal. Further, the vertical gradients of mean wind speed WU, mean
potential temperature § and mean specific humidity q are approximated by
the vertical differences 7, A@, Aq between these meteorological quantities
at a certain height and their pads at the sea surface. When doing so we
assume that at the sea surface the wind speed is equal to zero and the
temperature and moisture of the air are determined by the temperature of
the sea surface (saturation assumed) . If finally the wind stress tT is
replaced by the expression 1=pC_u 2 we arrive at the bulk
aerodynamic formulae Po

H = c,pC (8, - 6 )u and E = pC _(q_-q_)u (S)

where the subscripts O and a refer to the sea surface and the level
Z= a, respectively.

We see that the validity of this approach depends partly on the
applicability of the drag coefficient C, and partly on how well the
vertical gradients can be approximated by the vertical differences air-sea.
The first condition has already been dealt with and we are very well aware
of the difficulties and limitations inherent in the Cg concept. But what
is less well known,is the validity of the second assumption.

Therefore, I am going to examine this case more in detail. Some
suitable diagrams have been published by Brocks (1963). Profile measure-
ments of wind speed, air temperature, and water vapor pressure were made
by means of a buoy well away from land and also without any disturbance
from the accompanying vessel. The data used are averages each over a
period of 15 minutes.

Figure 8 shows the vertical wind shear between the 10 m and the 1m
levels as a function of the wind speed measured at 10 m. On the whole

Urg Uy ' H H i
m/sec ‘ °
°

Windprotile

Ostsee, Herbst 1958
° Vartihala Windscherung Uj,9-u, als
3 °5 oe Funktion des Windgeschraind ight Ugg
2) & Tae ot aT<-03°C 6
* 03°6 ATS 02°C ©
e © 02’< AT °
0 4 2 3 be BS 6 z 6 9 40 4 AZ m/isec ug

Figure 8. Vertical wind shear uj9-u, between the 10 1 and 1 m levels as
a function of wind speed uj and grouped according to the temperature
difference air-sea, AT. (from Brocks, 1963)

there is a clear relation between wind shear and wind speed but I am not

sure that it is a linear one. In addition, there is also a substantial
scatter caused by the thermal stratification as is indicated by the different
notations. Under stable conditions, the wind shear recorded with a certain
wind speed is distinctly greater than with adiabatic or unstable stratifica-
tion. Thus, for the height level z = a, we may write

du/dz ae £ (u,.A0_) (6)

if we try to be exact. Replacing the wind shear simply by the wind speed

as it is usually done, would introduce an error which depends on stability.

To avoid it, we must determine the function f(u,@,).
Similar relationships can be found for the potential temperature

and the humidity of the air. Figure 9 provides some evidence on the relation

between temperature gradient and temperature difference air-sea. Again a

definite - certainly non-linear - functional relationship is indicated,

Temperaturprofile
Ostsee, Herbst 1958
Vertikale Ditferenzen der potentiellen
Temperatur6,-8,als Funktion der pot.
Temperaturdifferenz 48 Luit-Wasser

Gruppen der ¢ Ou, 53 m/sec

o 3<u, 6 6 misec
° 6<u, 68 misec
d8<u,

Windgeschwindigkeit :

-50 -40 -3) -20 ~10 00 10°C 40

Figure 9. Vertical differences in potential air temperature NO ~ a
between the 10 m and 1 m levels as a function of potential temperature

difference air-sea 4® and grou according to the wind speed uw), measured
at 4 m height (from Brae TORS . : eS me

the scatter being caused by the wind speed as can be very distinctly seen
in the stable region. Consequently, we would have

98/a2| = £(ae_, u,) (7)

Finally, when looking at the results obtained from vertical profiles of
water vapor pressure (Figure 10) we recognize that a relation between

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ide! Wasserdamptprofile | Sper lies
omg Osizes, » April und eSept-Oht.1958 | : ee mee
| ae es

Verlikola Wosserdampiditterenzen €5-@,

rY) als Funktion von Bf y°e,-Ey Lmmbig] ia aia ae ae

- 02

-06

= e
06 ee °°

- 40

Figure 10. Vertical differences in water vapor pressure 210-e] between
the 10 m and 1 m levels as a function of water vapor pressure difference
air-sea Ae. The results of two research trips are distinguished by
different symbols (open and full circles). (from Brock, 1963)

gradient and difference air-sea apparently exists also here, but the
scatter is rather large, because - in addition to the humidity difference
air-sea - two other parameters - wind speed and stability - or dynamic and
convective turbulence - determine the moisture exchange occurring in the
boundary layer air-sea. In this case, we must write

de/dz ae f (de, u_» A8_) (8)

’ After inserting the empirical functions fi, fo, f, into our initial equations
(5) for the fluxes of sensible and latent heat we obtain

cy a ane
Co) Cc. f, (A0_ u,) us
fy (u,> Ae.)
Sant = =o (9)
sa ) Cc. f, (Aq_> u> Ae.) us
5 Cu,» Ae.)

These relations are considerably more complicated than the bulk aerodynamic
formulae normally used and render themselves not so easily to practical
application as the original relations do. They clearly show however the

deficiencies inherent in the usual bulk aerodynamic approach. Even if we
presume as a first approximation that the f-functions are linear in their
main variable, we would not get rather much farther. Then we would have the
following equations

(10)

ina]
iT}
no)
OQ
~
|
1
+O |
~
e|
Fh
~
S
ww
i= 4
@D
~

which contain two other empirical functions and decisively differ from the
original ones and from each other regarding their dependence on Ua, and
Ae.

Thus the bulk aerodynamic approach does no longer look as simple and
as adequate for practical use as before. In order to get a quick check,
let us assume that Brocks' graphs may be generally used and that further it
is sufficient to approach each of them by one linear relationship in each
ease thus disregarding the other influences. Even with such a simplification
we would obtain rather different coefficients for H and E, namely

=x
'

= 0.584 OL MO. (Os 0.) F

as hac laa (11)
E = 0.841 pee (qu ="qu)) u

This example shall only show what difficulties are implied when using the
simple bulk aerodynamic formulae. Certainly, the differences mentioned

might be compensated by a proper determination of C_. Therefore, great
care should be taken in assessing a particular value of C_ separately for
each of the two flux equations and for every special situation. Then a
rather high accuracy can be expected from this approach as was shown by
Webb (1960). Moreover, the concept of the so-called "profile coefficients"
(which are simplified forms of the functions f, fo, f3) may be used here
with some success.

After this critical treatment of the bulk aerodynamic approach let me
pass to some problems occurring immediately at the sea surface. While the
molecular viscosity does not play a decisive part, if any, in the exchange
of momentum at the sea surface, things seem to be different with the transfer
of heat and moisture. The momentum exchange is brought about by pressure
forces acting on the roughness elements. However, this model cannot be
applied to the exchange of sensible and latent heat, because, in this case,
there is no equivalent to the pressure exerted on the waves.

Thus we are led to believe that the final exchange of heat and
moisture between air and sea can only be effected by molecular processes.
Consequently, an adequate representation of evaporation and heat transfer
must include the molecular coefficients of conductivity and diffusivity.

A suitable approach has been described by Sheppard (1958) who, in the basic
exchange formulae, simply added the molecular constant v. and D to the
turbulent coefficients, Ky and Ky.

36
H = - =
@ © Ui sh _)) 5

oq. (12)
E=- (kK, +d) 32

Thus, there is no longer assumed a distinct layer of exclusively molecular
transfer; molecular and turbulent exchanges are rather supposed to exist
simultaneously. Since the turbulent coefficients decrease when the boundary
is approached, the molecular constants will become important very close to
the sea surface. I used this model with fair success for the computation of
marine evaporation on the basis of profile data measured close to the sea
surface.

Another phenomenon which is connected with the sea surface is the
existence of the so-called "cool skin," a-e-, a very shallow surface layer
where the water temperature is about 0.5 C lower than measured with con-
ventional methods which refer to a deeper layer. This cool skin, which
originates Prom evaporation, may be of importance for the thermal inter-
course between air and sea, since it is the actual sea surface temperature
that determines the character of the convective and turbulent motions in
the marine atmosphere. The observational evidence - given in part by small
sensing elements, in part by radiometers which, depending on their optical
properties, measure the water temperature of the uppermost layer of, say

1/10 mm - is somewhat varying. This may be explained by variable surface
film effects. It is well-known that surface films, i.e. monomolecular
layers consisting of organic matter or originating from artificial contamina-
tion can be found on the sea surface rather often. The transport of water
through a compressed and, therefore, oriented monolayer is not an ordinary
diffusion process which involves a small energy barrier but is to be con-
sidered as a process in which the water molecule must pass along a molecular
pathway between long molecule chains, thus requiring a substantially higher
amount of energy (La Mer, 1962). Consequently, a compressed monolayer
results in a retardation of evaporation. Laboratory studies have shown
(Haussler, 1955-56) that a cool skin does not occur in the presence of an
oil film, which prevents evaporation but does not impede the transfer of
heat. Thus it seems clear that the cool skin is actually due to evapora-
tion and that it will only exist on those parts of the sea surface which

are not covered by a monolayer. Hasse's (1963) paper provides further
evidence for this.

Quite recently it has been reported (Jarvis, 1963) that monomolecular
films not only retard evaporation but will also change the temperature of
the sea surface, because of the effect of the film on the convective move -
ment of surface water. The surface is constrained by such a film, convection
is found to be inhibited with a consequent thickening of the surface layer
and a rise of the surface temperature which is said to be quite independent
of any reduction of evaporation.

Field measurements indicating the reduction of evaporation by natural
surface films have been reported by Deardorff (1961). He compared the
evaporation rates of two pans floating at the sea surface. One of these pans
was filled with subsurface sea water while special care was taken in filling
the second so that any surface film, which might have been present, would
have been retained. The result showed a distinct reduction of evaporation,
of the order of 20 percent, which was obviously caused by natural surface
films. This surface film effect adds another item to the substantial list
of difficulties encountered when trying to measure marine evaporation.

Only very little is known about the effect that sea spray might exert
on evaporation, apart from that this influence must be present at higher
wind velocities. It is indeed very difficult to get any quantitative informa-
tion on this subject because field measurements of evaporation taken at the
sea surface can only be made with light winds where there is no appreciable
sea spray and, on the other hand, the problem seems hardly approachable by
means Of theoretical studies. Perhaps, quantitative information can best
be expected from laboratory work. Relevant studies have been carried out
by Okuda and Hayami (1959) in a wind-water channel of about 20 m in length.
They observed the vertical distribution of the horizontal transport by
sprayed water drops and found that the water transport by spray at 10 cm
height was negligible below 10 m/sec wind speed but increased rapidly in
the wind speed range of 12-13 m/sec. A strong decrease is found with
height, the water transport by spray at 30 cm being only about 10 percent
of its value observed at 10 cm height. The influence of sea spray on the

DIU

vertical moisture transport was studied in terms of the so-called “profile
coefficient," a number which is proportional to this exchange. As is shown
in Figure 11, a distinct increase of this profile coefficient is observed
for wind speed about 15 m/sec which is accompanied by a corresponding
increase of evaporation.

@ Nordsee 1959
@ Ostsee Herbst 1956
© Vindkanol (Okuda und Hayomi 1959)

i) 5 0 5 misec Uym

Figure 11. Profile coefficients I, of water vapor pressure at 4 m height
as a function of the wind speed u), measured at 4 m.
Full circles: Open sea (Brocks, 1959)
Open circles: Wind-water tunnel (Okuda and Hayami, 1959). (from
Brocks, 1963)

Of course, it would be very welcome to get some evidence on the sea
spray effect provided also by measurements in the field. Such measurements
are of interest not only because of the expected effect of sea spray on the
evaporation and vertical moisture transport, but also because it has been
suggested that the larger droplets ejected from the sea surface will be
accelerated by the wind and thereafter, when falling back into the sea, may
contribute to the downward transport of horizontal momentum, i.e. to the
wind stress acting on the sea surface. Relevant measurements have been
undertaken quite recently during the Aruba Expedition of the Woods Hole
Oceanographic Institution under the supervision of E. B. Kraus (1964). A
second expedition is now under way. There is some hope that some useful
results will be brought out with regard to the almost unknown effect of sea
spray.

After the critical remarks I made with regard to the bulk serodynamic
method some suggestions would perhaps have been appropriate on how the
vertical fluxes of sensible and latent heat could be estimated better.
Certainly, there are methods that can be applied with considerably higher
reliability, but they require a much greater instrumental expenditure.
Measuring the fluctuations would be the most direct approach but also

profile measurements (at three levels) would furnish acceptable values.
Finally, if these two possibilities cannot be used, some help can be
expected from those ‘profile coefficients" which enter the refined bulk
aerodynamic equations and which.can be obtained from some general tabulations
if wind speed and stability are known. It is, however. not possible here

to give more details on these procedures.

CONCLUDING REMARKS

If some final conclusion shall be drawn after reviewing the present
state of our knowledge on the physical exchange processes occurring at the
sea surface, then it should perhaps be concerned with the question whether
the relationships available at present for computing the vertical fluxes of
momentum, heat, and moisture passing through the boundary layer air-sea
might be serviceable and reliable enough for calculating these fluxes for
large sea areas on a routine basis so that synoptic charts of air-sea
exchange can be drawn in order to supply the necessary data for diverse
meteorological and oceanographic forecasting purposes.

The bulk aerodynamic equations certainly are easy to apply, and their
application can be based on such data as are normally available from the sea
at present. But it seems questionable whether they equally give what we
would like to have with sufficient precision. They appear to supply reliable
results if the coefficients inherent can be determined separately for each
quantity and also for each situation. However, I have some doubt whether
this may be possible if a general and permanent application has to be made
for a large sea area. Thus. the bulk aerodynamic formulae can only be
considered as a very preliminary and incomplete procedure for application
on & synoptic scale.

Therefore, we must look for other sources of information on the
turbulent vertical transfer quantities. The direct measurement by applying
the eddy correlation method certainly is not manageable generally because of
the complicated instrumental expenditure needed. However, the profile method
should be applicable in a general way. It would necessitate measurements of
wind speed, air temperature and humidity each carried out at three levels,
but it would be much better if their average vertical gradients at one
level (say 5 m) could be directly recorded with sufficient accuracy. This
object, of course, implies a considerable amount of development in instrumen-
tation and engineering but, in principle, it should be possible. Using such
devices we would get rid of all the problems connected with the drag co-
efficient and with the replacing of vertical gradients by differences air-
sea.

Sometimes I~@m dream@g of air-sea interaction buoys distributed over
the oceans in a fairly dense network and broadcasting regularly their
measured data on the mean vertical gradients which would enable us to draw
reliable charts on air-sea interchange. Let us hope that this dream will
not always remain a dream but will become true inynot too distant @ future.

Ashburn, E.V.

Benton, G.S., Fleagle, R.G.,
Leipper,D.F.,Montgomery,R.B.,
Rakestraw,N., Richardson W.S..
Riehl,H., and Snodgress,J

Brocks, K.

Brocks, K. and Hasse, L

Bruce, J.P., Anderson,D.V.,
and Rodgers, G.K.

Charnock, H

Clarke, D.B.

Darbyshire, J. and
Darbyshire, M.

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Some Results Obtained for Light Winds.
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A SURVEY OF THE ROLE OF SEA-AIR INTERACTION IN
TROPICAL METEOROLOGY

Joanne Simpson
U. S. Weather Bureau, Washington, D. C.

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Sea-air interaction affects every scale of motion and nearly every
process in the tropical atmosphere. It enters every problem in which we
attempt to make an explanation or a prediction - from interpreting the cloud
forms we see on satellite pictures to the oceanic semi-diurnal convection
eycle and the formation of hurricanes.

But what do we mean by sea-air interaction? How do the sea and air
affect each other in the tropics? Since we are mainly discussing meteorology
here and not physical oceanography or biology, let us examine more specifically
how the physical interaction between the sea and the air affects the atmos-
phere. This takes place by means of exchanging of property: primarily
moisture, heat, momentum and salt. Concerning each of these exchanges we
must ask three main questions, namely:

1. What is its magnitude and distribution in space and time?

2. What is it controlled by, that is, what is its functional
dependence?

3. What is its role in atmospheric processes and how important
in this role?

Momentum flux is intimately coupled with the others, as is pointed out
in the article by Roll in these Proceedings. Salt flux is probably important
in atmospheric thermodynamics as well as in the condensation process;
important research is going forward with its highly necessary documentation
(ef. Woodcock, 1958). This article, however, will restrict its subject to
moisture and heat fluxes, because of their large role in the energetics of
air circulations.

First, we will examine their global distribution and role in planetary
flows. This will lead naturally to a discussion of convection and cloud
patterning and the role of exchange in these processes. Finally, we shall
conclude with a more detailed consideration of the interaction between
oceanic fluxes and air circulations in the trade-wind and equatorial trough
zones of the tropics.

Knowledge of the role of sea-air exchange in atmospheric circulations
has burgeoned since World War II - in some cases it has even been incorporated
into models, both theoretical and numerical. There is, however, one major
reservation which relates to the most serious bottleneck facing sea-air
interaction studies and facing tropical meteorology - that is, for all
practical purposes, we have no direct way to evaluate these fluxes. We can-
not chart them from direct measurements, as we can sea temperature, for
example.

Nearly all our knowledge of heat and moisture fluxes from sea to air
is based on indirect calculations from the so-called transfer formulas or
Jacobs formulas. These are partly empirical and partly based on a simplified

model of turbulent boundary layer processes occurring at the sea-air
interface. This modeling should be best applicable when the windspeed is
high and the static stability is near neutral - that is, for near-zero
values of the Richardson number. The formulation gives increasingly bad
results as the Richardson number increases, to either positive (stable) or
negative (unstable) values.

This is not the place to develop the transfer formulas nor to discuss
further their range of validity, although a lot of work and consideration has
been recently devoted to this topic (Roll, 1965; Garstang, 1964). The fact
remains that flux computations based upon them, good or bad, form one of the
primary cornerstones upon which tropical meteorology has been built. Spot
checks of these flux computations have been made, by methods which are
themselves indirect and subject to errors and assumptions. Two main ways
of checking are the energy budget method, which involves assessment of
radiative fluxes, and the so-called "direct" method by aircraft measurements
made at some height above the surface (Malkus, 1962). These results nearly
always agree with those of the transfer formulas within a factor of two,
and often to better than 25 percent. It is all very well to reiterate the
truism that these fluxes just must be more accurately specified, to go
forward with tropical meteorology, but no one has yet produced a way to do
this, particularly on the necessary routine and frequent basis over wide
expanse of ocean.

The second important point to make here about the Jacobs transfer
formulas is that, to the extent that they are valid, they tell us that the
atmosphere itself mainly controls the extent of its own heat and moisture
input from the sea. We see this in the form of the equations, which is

Flux = Coefficient x (Air-Sea Property Difference) x Windspeed (1)

Since time fluctuations in the air-sea property difference are mainly
governed by those in the lower air, we can see that the atmosphere opens
and closes its own fuel line, making a very intriguing feed-back linkage;
the fine beginning made by Kraus (1959) in modeling this has not been
pursued as it deserves.

The coefficient in the transfer equations needs consideration prior to
interpretation of any of its results. In most climatological maps of heat
and moisture exchange, this coefficient is used as a constant. In his
classical work, Jacobs (1951) obtained his constant coefficient by
"calibration" with the energy budget method. There are more sophisticated
but probably no physically sounder ways of evaluating the coefficient
nowadays.

Boundary layer modeling indicates that the coefficient is a function
of Cp, the so-called "drag coefficient," which relates momentum exchange at
the anteerate to the windspeed. C., must be empirically determined.
Recently a number of workers have examined its dependence upon the atmospheric
variables (cf. Garstang 1964; Roll, 1965; Deacon and Webb, 1962; Deacon,

Sheppard and Webb, 1956; Sheppard, 1958). It is found to increase with
windspeed and with decreasing Richardson number - that is, the drag of the
sea on the air appears to be greater at lower stability with the same
windspeed. The drag coefficient roughly doubles as the wind increases from
2 to 14 meters per sec. Normal tropical variations in Richardson number
contribute only about one-fourth this much variation.

Keeping these problems in mind, we turn to our first topic, namely
the role of tropical sea-air fluxes in the planetary circulations. To
examine this, we need the climatological picture of these fluxes - the
magnitude of evaporation and sensible heat exchange - on annual global
maps, and analysis of regional and seasonal variations.

The classical maps of Jacobs (1951) are shown again in Figures 1 and
2. Malkus (1962) and Garstang (1964) have compared these distributions with
the later results of Budyko (1956) who also used the transfer formulas but
does not divulge his data sources or method of analysis. In looking at
these charts, we must keep in mind one further important limitation and
that is the data problem. Even supposing that the transfer equations were
exactly correct with an exactly known coefficient, for good results we
would need a measurement network of air temperature and humidity, sea
temperature and windspeed reporting every few hours. Then we should make
the multiplications required by equation (1) from each set of data and
average these products over month, season or year.

Of course, this is a visionary goal. Climatological mean values have
to be plugged directly into the formulas and clearly this can lead to errors
if there are correlations between air-sea property difference and windspeed -
if, for example, the air temperature commonly drops in storm situations.

This points the finger directly at synoptic disturbances.

Malkus (1962) and Garstang (1964) have made several case studies where
the transfer calculation from fairly long-period means could be compared
with averages of frequent measurements of the input into equation (1) from
research vessels or Weather Ships. Suffice it to say here that in the strong
and steady trades the correlation is unimportant, but wherever disturbances
are predominant, such as in the equatorial trough, the error can become
quite large. Garstang's (1964) results indicate that there it may be
considerably larger than the earlier estimates by Malkus (1962).

Figures 1 and 2 show by and large the expected flux distributions,
with some curious features, such as higher transfers in the Pacific than
in the Atlantic, which exhibits negative sensible heat flux off North
Africa. Budyko's (1956) values are greater in the Atlantic, suggesting that
another scale of fluctuations has perhaps distorted the picture. Jacob's
calculations were made using values at 1200 GCT, which is near midday in
the Atlantic area and at night in the Pacific. We do not know what observa-
tions Budyko used, but the diurnal transfer cycle that we shall describe
later suggests that he may have used these at 0000 GCT. In all transfer

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a meaningful interpretation of results.

Figure 3 shows the evaporation integrated by latitude and broken down
by seasons. Water vapor constitutes more than three-fourths of the
atmosphere's energy source; this is provided by evaporating an average of
just over 1 meter of sea water per year; about 75 percent enters the air
equatorward of 30° latitude. Figure 4 shors the oceanic heat budget calculated
by Budyko (1956), with roughly the same distribution of evaporative heat loss
(Q. = LE, where L is the latent heat of vaporization) as shown in Figure 3
from Jacobs. Note the very mch smaller magnitude of Q., the sensible heat
supply from sea to air; the ratio of Qs to Q, however, rises outside the
tropics, for reasons that are discussed later.

Budyko's Q, shows a similar equatorward dip as Jacob's does, a deduction
very crucial to the oceanographer as well as to the meteorologist. Q,, is
the oceanic heat flux divergence. In most studies, including Budyko's, this
is deduced as a residual in the oceanic heat budget; it is essentially the
difference between the radiation balance R and the heat loss by evaporation,
Qe. This heat energy difference is what the equatorial ocean has left to
export to high-latitudes, to moderate their winter climates through warm
currents such as Gulf Stream and Kuroshio. By integrating Ova latitudinally,
authors like Bryan (1962) use the only way extant to arrive at oceanic heat
transports; these suggest that the oceans may carry as much as 15 - 20 percent
of the heat energy transported by the general circulation of the atmosphere.
Either this important result, or global radiation figures, must suffer an
agonizing reappraisal if Qe im equatorial regions undergoes significant
alteration, as Garstang's contribution to these Proceedings suggests it must.

How does the air utilize these energy inputs from the oceans? To
begin to answer this question, let us examine Figure 5, the heat budget of
the atmosphere. First note the latitudinal uniformity of the radiation
sink R,, which corresponds to a cooling of about 0.75°C per day. From 60°%N
to 60°S, neither the sensible heating from the surface Q,, nor heat flux
convergence in the air (-Qya_) does much to make up this large radiational
deficit - in fact the large negative peak in -Q,, (air heat flux convergence)
near the equator only compounds the air's heat losses in the tropics. As
we see from the top curve in the diagram, LP (precipitation heating), is
the vital atmospheric heat sourge which makes up both the Vas Loss
and provides for the heat export from the equatorial zone «4

Clearly the air must have converted the latent heat gained from
evaporation into the usable or sensible form by making rain. Comparison
of Figures 4 and 5 show clearly that the main input and the main utilization
of the water vapor occur in quite different regions - our attention is
directed to the wind circulations and cloud formation process to explain
this difference.

i] In latitudes poleward of 60°, heat flux convergence becomes as
large or larger than precipitation warming.

110

100

North Latitude
30°

Figure 3 (after Jacobs, 195la). Integrated latitudinal dependence of
Svevoza tien from N. Atlantic and N. Pacific Oceans together, by seasons. Units
109 m 3/aay. Mean annual evaporative loss to both oceans: 112.5 cm year-1,

TH
Mean Annual Heat Budget

of the Oceans

(after Budyko)
120

-20

N S

Degrees Latitude
Figure 4 (after Budyko, 1956). Annual heat budget of the ocean. Units,
kg calories per cm“ per year. Q, is evaporation; Q, is sensible heat
transfer; R is the radiation balance, and Qy, is the oceanic flux
divergence, computed as residual to balance the budget.

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Figure 5 (after Budyko, 1956). Annual heat budget of the atmosphere. Units,
kg calories per em“ per year. LP is precipitation warming where P is
precipitation in cm and L is the latent heat of vaporization. Qs, is the
sensible heat transfer from the ocean; - Qya is atmospheric flux convergence

of sensible heat plus potential energy, and Rg is the radiation balance of
the atmosphere.

20 40 60 5s

In Figure 6, low-level wind patterns begin to delineate the intriguing
paradox: equatorward - blowing trades carry the water vapor fuel unconverted
into the convergent trough zone (solid line) - it is largely not processed
locally in the input zone and from there exported poleward, but goes on
this roundabout circuit because of the air's movements and differing cloud-
building abilities.

Trade-wind clouds are normally stunted cumuli, as depicted in Figure 7,
while in the equatorial regions, high cumulonimbus towers (Figure 8) often
flourish. The latter are excellent latent heat converters and not the former
which usually evaporate without dropping precipitation back into the ocean,
which is the prerequisite that the heat stay in the air.

We see that the cloud and precipitation processes control the role of
sea-air exchange products in the atmosphere - so again we return to air
structure and motions and this time ask how these interact with clouds. In
1957 we made an air»orne photogrammetic study in the Pacific to begin an
attack on this question; the complete results have just been published
(Malkus and Riehl, 1964). From the carefully time-lapsed movies, cloud maps
were constructed and compared with sounding and synoptic data. Figures 9 and
10 show two typical trade-wind cases, the first with weak winds and the
second with normally vigorous flow. In both, the ocean is presumed but not
known to be warmer than the low-level air. To assess the role of sea-air
interaction, comparison with Avsec's (1939) extension of Benard's classical
convection cell studies was made. Figure 11 shows one of Avsec's laboratory
experiments, with uniform heating from below and weak translation (weak
shear between convecting fluid and lower boundary). Figure 9 suggests a
similar admixture of polygonal cells with cells becoming elongated into rolls.
In Figure 10, the stronger flow has presumably caused the rolls to predominate;
our Pacific study confirmed the conditions, related to shear, for their
orientation.

From satellite pictures, Krueger and Fritz (1961) have identified
polygonal cells in some way apparently similar to those studied in the
laboratory. Figure 12 is from one of the few cases where they had enough
supplementary data to relate these patterns partially to sea-air interaction.
The sea-air temperature excess was large ( ~ 3°c ) and a strong trade-wind
inversion confined the convection - its horizontal scale increased toward the
southeast as the moist layer deepended. Three puzzling features were found,
however, prohibiting the naive extrapolation of laboratory results.

Firstly, the winds were strong (15-20 knots) in the region (Figure 13)
where the polygonal pattern was photographed - why were not the polygons
elongated into rolls? Secondly, the horizontal cell size runs about 30 times
the depth of the convective layer, exceeding the laboratory ratio by an order
of magnitude. Thirdly, Figure 12 (upper) suggests each cell wall contains
several cloud elements, or more than one scale of motion.

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polygonal cells into longitudinal bands, by the setting
in motion of a layer of air heated from below. Depth
of layer 20 mm. Compare with cloud map of Figure 9.

Figure 12 (after Krueger and Fritz, 1961). Case of polygonal cloud
cells formed over subtropical Atlantic Ocean on April 4, 1960.

Narrow angle picture (upper) corresponding to wide angle

picture (lower). This outlined area is approximately 100
nautical miles square.

(after Krueger and Fritz, 1961). Surface weather map for

1800 GCT, April 4, 1960. Track of Tiros I, subsatellite point
(dot), and picture center (circled dot) are indicated. The
area occupied by cellular cloud pattern in Figure 12 is out-
lined by dashed line.

Clearly the fact the atmosphere is fully turbulent complicates the
development of quantitative criteria for cloud patterning and makes difficult
the assessment of the role of sea-air interaction in modifying or regulating
these patterns. However, we can learn much more about these relations if
we would only focus our existing observational resources simultaneously upon
the sea-air boundary conditions, the air structure, and the cloud patterns.
We do not yet know the relative roles of the first two in producing the
fascinating cloud configurations that the satellites are showing us. For
example, in Figure 14 as we move (from left to right) following the southeast
trades toward the equator, the cloud patterns change from cellular to actino-
form (Picture of the Month, 1963) to blob-like. Hubert 1/ has postulated
that in going from cellular to actinoform, heating from below may be giving
way to dominant cooling from above. The change to blob form in the equatorial
zone may be due to deepening of the convective layer. Wo data exist to test
these interesting and important suggestions. The potential of satellite
pictures as a tool to understand and predict tropical weather, to interpret
the role of sea-air interaction therein, cannot progress much farther until
ship and aircraft observations are jointly focussed on these situations. The
large necessary expenditure dwindles when compared to that of the space
program.

A problem to isolate for the next step of the investigation might be
the abrupt transitions in oceanic cloud forms. A typical example is show
in Figure 15. Are these common transitions, often almost infinitely sharp,
related to sudden changes in sea temperature or are they governed by swift
transitions in air motions? We know they are more abrupt than any likely
structural changes in the atmospheric sounding, in terms of lapse rate or
humidity stratification.

The beginnings of the required study were made with the Woods Hole
Aircraft (Malkus, 1957) using an airborne radiometer, soundings, and cloud
photography. "Warm spots" in the ocean upper layers were found, which were
also documented by a research vessel with thermistors at several depths.
These warm spots were related to trade cumulus groups observationally
(Figure 16). The "heated island" and “equivalent thermal mountain" frame-
work (Stern and Malkus, 1953) was used to explain the observations. Even
these small-amplitude warm spots should produce an “equivalent thermal
mountain" about 100 meters in elevation, which can be quite effective in
triggering trade-wind clouds. In continuing these studies, a difficulty is
that the airborne radiometer must be flow below 1000 ft., while cloud photo-
mapping is much better performed at 10,000 ft. or more. Furthermore, a
surface vessel is needed to calibrate the radiometer and to determine the
depth of any oceanic temperature anomalies found. Despite these complications,
such a program is mandatory to advance the interdependent fields of sea-air
interaction, satellite interpretation, and tropical meteorology.

i Personal communication to the writer

oud pattern.

Tiros V, 1500 GMT,

, 1963).

Tropical (southern hemisphere) cl

Figure 14 (after Picture of the Month
October 7, 1962

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Figure 16 suggests the close relationship between active sea-air fluxes
and cumulus convection. Although a temperature difference of 0.17°C may
appear very small, it amounts to 30-50 percent of the common sea-air
temperature difference in the trade-wind region. Therefore, we may presume
that, over these warm spots, the flux of latent and sensible heat from sea
to air is somewhat greater than in the surrounding regions, as are the
vertical fluxes of heat and moisture in the well-mixed layer that character-
izes the air below cloud base.

The important relation between variations in sea-air temperature dif-
ference, exchange and cumulus development is illustrated in the semi-diurnal
cycle in convection over the tropical oceans.

A midday minimum in oceanic cumulus convection was long suspected and
finally has been substantiated, primarily by the two cruises of Garstang
(1964) with the Research Vessel CRAWFORD. Oceanic cumli flourish best near
dawn and sunset and undergo a suppressed period during the noonday hours.

At least part of the cause appears to lie in a similar cycle of sensible
heat flux (Figure 17). There is only a weak variation in latent heat flux
(Figure 18). Figure 19 shows that the daily variation in sea-air temperature
difference provides the main key to the transfer cycles, although there is
also a weak wind minimum in the early afternoon, related by Lavoie (1963) to
the So component of atmospheric tide.

Interestingly enough, the convergence - divergence cycle of this tidal
wind is quite nicely in phase with the cloudiness ( Figure 20). Shibata
(1964) showed that the amplitude of the convergence (+ 2-3 x 107 sec! )
and its attendant vertical motions can be related quantitatively to the
cumulus variation. Which is cause and which effect? Does the semi-diurnal
cycle in air-sea heat transfer help to drive the atmospheric tide via the
cumalus process? Or does the tide, maintained solely in the high atmosphere,
add to the air-sea interaction cycle in maintaining that of the cumulus? A
very little more research could go a long way toward answering this intriguing
question.

A recurring thread is clearly woven throughout all these examinations
of the role played by exchange in tropical meteorology: This role is
directed by clouds. How the atmosphere uses what it receives from the sea
depends on what sort of clouds its circulation enables it to produce. And
this, in turn, depends on large-scale atmospheric dynamics, or the air
motions throughout the depth of the troposphere.

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energetics of the trade-wind zone and the equatorial trough. The amount
of sensible and latent heat that the sea puts into the air does not differ
highly between these two regions of the tropics. And yet, its usage is quite
different, as is the cloud structure.

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In the trade-wind stream, the cloud development ( Figure 21) is

restricted by the trade-wind inversion, so that only about 25 percent of
the latent heat gained is released by local precipitation. The remainder
is accumulated in a deepening moist layer and is shipped toward the equator
by the trades ( Figure 22). The rain release occurs at fairly low-levels,
and together with Q., provides a net heat source and “pressure head" which
sustains the lower trades. A preliminary mathematical model of this system
has been constructed (Malkus, 1956). The upper trades live on more sporadic
imports from the equatorial trough zone.

Figure 23 summarizes the energetics of the region and the role of
exchange therein. Sensible heat plus potential energy (h = Gof + Agz )
is budgeted on the left and latent heat L, (q is specific humidity ) on
the right. Three layers are treated separately; in ascending order. These
are roughly consistent with the mixed layer, cloud layer and above-inversion
layer (although 500 mb is generally above cloud tops). The important point
is that the sum of Qe + Qg (1-87 units) is easily enough to balance the total
radiation loss (1.32 units) but camnot do so, since 1.41 units of latent are
exported. Half of Q, balances the radiation loss below cloud, and the other
half, together with all the precipitation, does not quite balance the radia-
tion loss of the middle layer. The heat import at high levels makes up the
difference. The upward - directed dashed arrows show the important "heat
pump" function of the cumuli.

Convective clouds play an even more crucial role in the operation and
energetics of the equatorial zone; they provide the release mechanism for
most of the latent heat energy acquired from the sea over the entire tropics.
Here giant cumulonimbus towers abound (Figure 24) but intermittently, bunched
into the wavelike and vortical perturbations, with horizontal diversions
ranging from about 200 to 2000 km.

The heat budget for the equatorial zone is dissected in Figure 25.
Here we are examining total heat content ( Q = C,T + Agz +L, ) The input
from the ocean is computed as 70 percent that of the trade-wind belt. Qe.
was obtained from the transfer formla and Qs was found as residual in an
energy budget which may have exaggerated its magnitude, but not by a two
factor. The main feature is the high heat export aloft. In their detailed
study, Riehl and Malkus (1958) showed that the penetrative cloud towers were
necessary to get this heat energy upward, balancing radiation losses and
providing the crucial export.

In trade zone and trough together, the ocean provides a total Qs + Qe
of 3.18 units. Of this, 73 percent is eventually lost locally by radiation
from the tropical atmosphere, leaving only 27 percent of the heat energy to
ship poleward across the subtropical ridge. Only a few percent of this
heat is ever converted into motions. As well as an inefficient heat engine,
we see that the atmosphere also operates a very leaky fuel pump!

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100

The interaction between exchange and air circulation has been re-
emphasized in this last discussion. The patterns of exchange are patently
somewhat different in the equatorial trough zone from what they are in the
trades - their relations to convective clouds and cloud distributions are
also strikingly different.

The fluxes themselves are more variable in the trough zone than in
the steadier trades. Frequent disturbances dominate the scene here,
particularly in summer. By producing large clouds, these disturbances
provide the mechanism which make the main product of sea-air exchange
(moisture) usable in the atmosphere.

In the trough zone, another imprint of the disturbances is the much
increased Q,/Q, ratio. The important documentation of why and how this
ratio is higher in disturbed zones is taken up by Garstang in the following
contribution to these Proceedings. We merely conclude here with two
important questions which bridge the two topices; these are:

1. How do disturbances affect sea-air interaction in the tropics?

2. What role does sea-air interaction play in the life and growth
of disturbances?

Concerning the first question, the foregoing has already given a clear
indication that tropical disturbances are critical in regulating sea-air
exchange. Garstang's work illustrates this in a rather disturbing way, in
that the results suggest that we may have to make drastic changes in the
accepted climatological exchange picture. This could upset the global
budgets and part of oceanography as well.

In respect to the second question, we have firm evidence that in the
hurricane stage, local sea-air fluxes play a large and crucial role in the
machinery of the storm's heat engine (Malkus and Riehl, 1960). In sub-
hurricane disturbances, we do not know what the role of local exchange is,
nor do we yet have adequate models to form a framework for such a study.
Quantitative observational studies of the sort described in the following
paper by Garstang thus constitute the essential next step.

101
REFERENCES

Avsec. D. 1939: Thermoconvective eddies in air.
Application to meteorology. Scientific
and technical pub. of Air Ministry Work
of Inst. of Fluid Mech. of Fac. of Sci.,
Paris, No. 155. Published (in French)
at Ed. Blondel la Rougery-., 7, Rue
St. Lazare, Paris.

Bryan, K. 1962: Measurements of meridional heat
transports by ocean currents. Jour.
Geophys. Res., 67, 3403-3414.

Budyko, M. I. 1956: The heat balance of the earth's
surface. Gidrometeorologicheskoe
izdatel'stvo, Leningrad, 255 pp.
(Translated by Nina A. Stepanova;
translation distributed by U. S.
Weather Bureau, Washington,D.C.1958).

Colon, J. 1960: On the heat balance of the troposphere
and water body of the Caribbean Sea.
National Hurricane Res. Proj., Rep't
No. 41, U. S. Dept. of Commerce,
Washington, D.C. 65 pp.

Deacon, E. L.; P.A. Sheppard; 1956: Wind profiles over the sea and the
and E. K Webb drag at the sea surface. Austral. J.
Phys., 9, 511-541.

Deacon, E. L. and E. K. Webb 1962: Interchange of properties between sea
and air. Small-scale interactions.
Chap 3, The Sea, Vol., 1, 43-87.
Interscience Publishers, New York and
London.

Garstang, M. 1964: The distribution and mechanism of
energy exchange between the tropical

oceans and atmosphere. Phd Disser-
tation, Dept. of Meteorology, The
Florida State University, Tallahassee,
Florida. 177 pp-

Jacobs, W. C. 1951: Large-scale aspects of energy trans-
formations over the oceans. Compendium
of Meteorology, Amer. Met. Soc.,

Jacobs, W. C. 195l1a:The energy exchange between the sea
and the atmosphere and some of its
consequences. Bull. Scripps Inst.
Oceanog., Univ. of Calif., 6, 27-122.

102

Kraus, E. B. 1959: The evaporation-precipitation cycle
of the trades. Tellus, 11, 147-158.

Krueger, A.F. and S. Fritz 1961: Cellular cloud patterns revealed
by Tiros I. Tellus, 13, 1-7-

Lavoie, R. L. 1963: Some aspects of the meteorology of the
tropical Pacific viewed from an
mmeilil, Seale IED"Bo Wo Fp
Meteorology Div., Hawaii Inst. of
Geophysics, Univ. of Hawaii.

Malkus, J. S. 1956: On the maintenance of the trade winds.
Tellus, 8, 335-350.

Malkus, J. S. 1957: Trade cumulus cloud groups: Some
observations suggesting a mechanism of
their origin. Tellus, 9, 33-44.

Malkus, J. S. 1962: Interchange of properties between sea
and air. Large-scale interactions.
Chap. 4, The Sea, Vol., 1, 88-294.

Malkus, J. S. and H. Riehl 1960: On the dynamics and energy trans-
formation in steady state hurricanes.
Tellus, 12, 1-20.

Malkus, J. S. and H. Riehl 1964: Cloud structure and distributions
over the tropical Pacific Ocean.
Univ. of Calif. Press, Berkeley,
Calif., 229 pp.

Picture of the Month 1963: Monthly Weather Review, 91, 2.

Riehl, H. and J. S. Malkus 1958: On the heat balance in the equatorial
trough zone. Geophysica (Helsinki),
6 (3-4), 503-538.

jomal Selo le 1965: On the present state of knowledge on
air-sea boundary layer problems.
Proc. Sea-Air Interaction Conf.,
Tallahassee, Florida, Feb. 1965,pp 31-63.

Sheppard, P. A. 1958: Transfer across the earth's surface
and through the air above. Q. J. Roy.
Met. Soc., 84, 205-22h.

Shibata, E.

Stern, M. E. and J. S. Malkus

U. S. Weather Bureau

Woodcock, A. H

Wyman, J. et al.

1964:

1953)

1938:

1958:

1946:

103

The atmospheric tide hypothesis on
the diurnal variation of cloudiness
in the tropics. M. A. dissertation,
Dept. of Meteorology, Univ. of Calif.
(Los Angeles), 82 pp.

The flow of a stable atmosphere over
a heated island, II. J. Meteor.,
10, 105-120.

Atlas of Climatic Charts of the Oceans.
Gov't. Printing Office, Washington,

D. C., Weather Bureau No. 1247,

130 pp.

The release of latent heat in tropical
storms due to the fall-out of sea-salt
particles. Tellus, 10, 355-371.

Vertical motion and exchange of heat
and water between the air and sea

in the region of the trades.
Unpublished report on file at The
Woods Hole Oceanographic Institute,
Woods Hole, Mass.

Re ee

oat faut

vty Moma
Ne ER iy iis

SENSIBLE AND LATENT HEAT EXCHANGE IN LOW

LATITUDE SYNOPTIC SCALE SYSTEMS

Michael Garstang
Florida State University
Tallahassee, Florida

105

106

ABSTRACT

The transfer equations are applied to a set of 46 days of data col-
lected from an oceanographic research vessel on a station in the low latitude
western Atlantic. In this region the lapse rate and shear in the boundary
layer of the atmosphere is such that the transfer equations may be applied
with a reasonable degree of confidence. In particular, the most accurate
results are likely when large transfers occur, the least accurate when small
transfers occur. Under these conditions a linear dependence of the eddy
exchange coefficient with wind speed and stability were used to compute
latent and sensible heat transfer from time averages over 1 hour of specific
humidity, temperature and wind speed at 6 m and at the sea surface. Sea
surface temperatures were measured at 10 cm and compared with infrared
radiometer measurements.

Individual synoptic scale systems that moved over or close to the
point of observation are examined. Over limited regions of these disturbances
latent and sensible heat transfers are found to increase by an order of
magnitude. Integrated over the whole disturbance the energy flux is found
to be double the undisturbed values. By using both streamline analysis and
a rainfall amount and occurrence technique, the frequency and size of synoptic
systems are determined. This makes possible the construction of summer,
winter and annual maps of latent and sensible heat transfer for the tropical
Atlantic. Significant differences are found when compared with results of
earlier workers. The role of synoptic scale disturbances in the atmospheric
energy balance is emphasized by these results.

107

I. INTRODUCTION

Considerations of the energy sources, conversions and transports in
the earth-atmosphere system soon lead to the conclusion that the tropical
oceans play an extremely important role in the heat and energy balance of
the atmosphere. The greater part of the incoming short wave solar radiation
in either direct or diffuse form is absorbed at the earth's surface. Before
this energy can be utilized by the atmosphere it must be transferred across
the earth-atmosphere interface. Therefore, the atmosphere is fuelled mainly
from below. As show by Malkus (1962, p. 93) and others, the most important
transfer process at the earth-atmosphere interface is evaporation which
provides about half of the atmosphere's fuel in the latent form of water
vapor. More than half of this water vapor fuel is supplied to the lower
troposphere by the tropical oceans between 30°N and 30°S latitude. If we
add to these considerations the observation that the overall radiation
balance of the earth-atmosphere system is positive between about 35°N and

fo)
35 S latitude and negative poleward, the validity of the opening statement
is justified in general terms.

In the tropical regions of heat input the net transfer of water vapor
and sensible heat is from the ocean to the atmosphere. Shear turbulence
within the first tens of meters of the trades insures thorough mixing and
upward transport of both water vapor and sensible heat. Turbulent eddies
continue the upward transport through the subcloud layer to the cloud layer
where convective cells, in the form of cumulus clouds, probably play the
dominant role in the vertical transport of energy. Within the undisturbed
trades high values of wind steadiness occur through a considerable depth of
the lower troposphere. Under these undisturbed conditions, gradients are
likely to stabilize and the energy input at the surface will reach an
equilibrium value dependent upon the vertical removal from the surface and
upward transportation of energy. Riehl, Malkus and others (1951) demonstrated
that these processes create a moist convective layer gradually deepening
along the airflow. In the trade wind regions of high evaporation most of
the moisture is retained in vapor form rather than rained back into the
oceans. This accumulated water vapor is transported equatorward “at a rate
of energy export easily two orders of magnitude greater than the rate of
kinetic energy consumption by all the global winds and sea currents combined."
(Malkus, 1962.)

Within the equatorial trough region much of this water vapor is
condensed, releasing its latent heat, which is carried to great heights in
cumulonimbus clouds. Riehl and Malkus (1958) have shown that relatively few
such large clouds concentrated in a limited number of synoptic scale systems
(e.g., the vortical and wave-like pertubations common in this region) are
able to transport this energy, now in the form of sensible heat and potential
energy. But now a distinction should be made between suggestions based upon
mean budget conditions and deductions based upon day-to-day synoptic changes.
If average trade wind conditions prevail for some period of time, gradients

108

of water vapor and temperature will approach some constant value and the
magnitude of the energy transports in the trades will be governed by this.
Within the trade wind-equatorial trough system synoptic scale disturbances
will produce internal convergences of water vapor and sensible heat and a
vertical flux of energy. But synoptic scale systems, as was shown by the
CRAWFORD cruise of 1957 (Garstang, 1958), significantly increase the trans-
port of both latent and sensible heat from the ocean to the atmosphere.
Gradients, particularly of temperature, in the air immediately above the

sea surface steepen markedly and wind speeds increase during synoptic
disturbances. The total flux of water vapor and sensible heat at the surface
over the area of a synoptic disturbance is, therefore, likely to be signifi-
cantly greater than during undisturbed conditions. Thus the mean picture is
constrained by overall planetary budget requirements and departures from it
are primarily due to the travelling synoptic systems which. form a vital link
between the ocean and the atmosphere. An attempt will be made here to
evaluate the role that organized synoptic scale disturbances play in determin-
ing the distribution of latent and sensible heat transfer over the tropical
oceans.

II. ANALYTICAL TOOLS AND AVAILABLE DATA

In restricting our attention to the tropical oceans we encounter the
fortunate situation that the lower decameters of the tropical atmosphere are
by and large barotropic and close to neutral stratification. Under these
circumstances, the most practical equations for the computation of the flux
of latent and sensible heat are the bulk aerodynamic equations expressed
below:

"
L@)
"
ae)
&
ie)
-~-~
10 |
'
+O |
sa
e|

; ie nee ie (1)

vo)
i
no)
fo
io)
V—_
@D
'
}|
—_
5 |

s p D ° a Qo (2)

where Qa. and Q, are, respectively, the eddy vertical transport of latent

and sensible heat and are directly proportional to the differences between
the mean specific humidity and mean potential temperature measured at the
surface and at some point above the surface multiplied by the mean wind
speed measured at the same point above the surface. If the atmosphere above
the sea surface is close to neutral stratification, then the most restrictive
assumptions that must be made in order to arrive at the equations are most
nearly satisfied.

Forty-six days of observations were made on two separate cruises, each
of 23 days duration, on a fixed station. Both cruises took place during the
months of August and September. The first in 1957 when the ship (RY vi
CRAWFORD, Woods Hole Oceanographic Institution, Garstang (1958)) was
stationed near 11° 52°W; the second in 1963 when the same ship was stationed
near 13°N 55°w (La Seur and Garstang (1964)).

109
The degree to which conditions in the air above the water departed
from neutral stratification was assessed by computing hourly values for the
bulk Richardson Number, R3>
Ty) igs (3)
r2

g
RR. Sf
T 2

u.
a

where g is the acceleration of gravity; T is the air temperature in degrees

Kelvin at height a; z is the height of the observation (6.0 m) above the

sea surface; @, the potential temperature at 6.0 m3; @5 the potential

temperature at the sea surface; ua the wind speed at 6.0 m; each of the latter
quantities averaged over 1 hour centered on the hour. The coefficients [,

and TI represent the gradients of heat and momentum, and are defined as

1 du
heats ga we
a
and
il 60
Nea a (5)
Oe - OR 6 ln z

According to Deacon and Webb (1962) at heights of about 4.0 to 8.0 m and for
conditions from neutral to moderate thermal stratification, [f, and [I
have values around 0.1. Thus the ratio in (3) was taken as equal to 18.

Ninety-six percent of the observations obtained on the CRAWFORD fall
in the range

“0.325 < R, < +0.025

3
sixty percent fall in the range

-0.040 < R, < +0.020.

Conditions are, therefore, relatively close to neutral stratification. The
departure from neutral stratification is a combined effect of lapse rate

and wind shear. If we consider the bulk of the observations in the range
-0.325< Rp <+t0.025 and assume a mean wind speed of 5.0 m sec”, an air-sea
temperature difference ranging between -4.06C (sea warmer than air) and
+0.21C (air warmer than sea) can be predicted from equation (3). The actual
maximum range (except for three observations in showers) observed was -2.88¢C
to +0.17¢C . It is, therefore, concluded that the most critical parameter in
determining stability is the wind speed and, more important, the stronger

the wind, the more closely governing conditions are met.

110

The saving grace for formations such as (1) and (2) for the computa-
tion of sea-air energy transfers is that, by and large, the higher the
transfer, the more accurately these formulae predict it. As will be show
in later sections, the largest exchange occurs at times of strong winds.
Here the shearing is unlikely to permit maintenance of large lapse rates
over the open ocean and, even if these did occur, the strong wind shear keeps
the Richardson number in a reasonable range. When the wind approaches calm
and the air-sea temperature difference remains large (unlikely), the formulae
are in trouble and give particularly bad results for sensible heat exchange.
But all the exchanges are then relatively small and perhaps even negligible
for large scale budget and dynamic considerations.

The preceding establishes the fact that if the bulk aerodynamic
equations can be used at all, they can probably be best applied to conditions
prevailing over the tropical oceans. However, it is also quite clear that
the drag coefficient, Cp, is not a constant but is a function of, at least,
height (z); wind speed (u) and stability (Rg). By applying considerations
outlined by Monin and Obukhov (1954) it can be shown that

mE 3/2 / u.

(KS 3 CH ( J? = aR Ck ii Pt) 2 dat 2

6 l= pl Co ri
6 bar}

(6)

where C* is the drag coefficient for neutral or adiabatic conditions and
the subscripts 6 and 11 refer to heights (in meters) above the sea surface.
A linear dependence of the drag coefficient upon height and wind speed under
adiabatic conditions was assumed using values obtained by Deacon and Webb
(1962). In practice (6) was approximated and the contribution of the term
in brackets on the right hand side was neglected. Figure 1 shows the
functional relationship between the remaining terms of the above expression.
Based upon these considerations the drag coefficient was then computed from

e = -3
Ce = (1.46 + 0.07 u, 4.2 R,) x 10 . (7)

Finally, before using equations (1), (2) and (7) to compute the transfer of
latent and sensible heat, an evaluation of the accuracy of the measurements
must be made. Wet- and dry-bulb temperatures and wind speeds were measured
at 6.0 m above the sea surface at a point 4.7 m on a pulpit ahead of the
bow of the vessel. Sea surface temperatures were measured 10 cm below the
surface and 4.5 m ahead of the vessel. The temperature sensors were therm-
istors and the wind speed was obtained from a recording cup anemometer.
Care was taken to avoid radiation effects. Supplementary temperature
measurements were made around the ship (out to 140 ft) on three different
occasions. These measurements were used to determine any extraneous effects
on the pulpit observations. Wind speeds were checked for ship effect in a
manner similar to that described by Deacon et al. (1956). It is felt that

dui

+0.025 QO -0025 -0050 -075 -0O1 “25 -OIS <I75 -0.2 <225 -025 -275 -0.3

Roe %/r 2 Te Oa
Tz uv
Figure 1. Departures from the neutral value of the drag coefficient
~ at 6.0 m as a function of stability expressed in terms of

the bulk Richardson number.

with these precautions and the corrections later applied, the observations
are representative of ambient conditions over the open ocean. =

III. SYNOPTIC VARIATIONS OF LATENT AND
SENSIBLE HEAT TRANSPORT

Im order to examine the distributions of latent and sensible heat
transfer under varying synoptic scale conditions, each observing period
was subjected to careful analysis. The objective was twofold:

1. To delineate synoptic scale disturbances and categorize each
system with respect to intensity;

2. To relate the observations made by the ship to each system
and delineate the sector intensity that was sampled.

To achieve this, conventional methods of analysis were supplemented
by a number of other techniques. Among these were: divergence, vorticity
and vertical motion computations using the Bellamy (1949) triangle method
for each of the five triangles shown in Figure 2, rainfall analysis of six
islands in the Lesser Antilles using a classification based upon occurrence
and amount, and a method using wind speed and steadiness as a measure of the
occurrence and strength of a synoptic scale disturbance.

1/ A more complete treatment of ship effect is contained in a doctoral

~ dissertation(Garstang (1964)). Similarly, details of analysis and
computation referred to in subsequent sections of this paper may be
found in the same source.

112

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Once the systems had been identified, time distributions of latent
and sensible heat could be obtained by calculating the hourly values for
all 46 days and obtaining average values (Q and Q_) for each of the
systems. Table 1 shows the initial broadsc&le breakdown. In the first two
rows all undisturbed states can be compared with all disturbed states. The
results from both sets of observations show the same trend: more heat is
transferred from the ocean to the atmosphere under disturbed conditions than
during undisturbed conditions. The increase in.latent heat transfer is not
as large as in the case of sensible heat transfer. The ratio of Q
undisturbed versus Q. disturbed is 1.17 in 1957 and 1.07 in 1963; fut for
Q, it is 7-74 and 2.04, respectively. On each cruise weakly disturbed
modes were fairly frequent and strong trade wind conditions relatively few,
so that it is perhaps more meaningful to examine the transfers under weak
and moderate trades and moderately disturbed modes. Under these conditions
the ratio of undisturbed to disturbed for Qe increases to 1.94 in 1957
and 1.25 in 1963; amd for Qs 11.70 and 2.16, respectively. The relative
increase in the Bowen ratio remains essentially the same for both of the
above comparisons. While the amount of sensible heat made available to the
atmosphere never exceeds 8 percent of the latent heat transfer, it is
Significant to note that:

ale This is an average figure throughout the whole disturbed period;

2. For periods up to 6 hours within a disturbed region, the ratio
of sensible heat to latent heat may be greater than 0.20 as opposed to
undisturbed conditions when, for similar period, the sensible heat transfer
may be in the opposite direction, i.e., from the atmosphere to the ocean.

3. This energy is directly available to contribute to the instability
of the subcloud layer and the development of convective cloud.

Therefore, it is suggested that the role of sensible heat transfer in the
formative and developing stages of tropical disturbances is a critical one.

The synoptic scale fluctuations of exchanges are examined in more
detail in Table 2. The mean values for each regime or mode are based_upon
values computed for each hour through each state. Little change in 4q _
is noted between the various states. The mean temperature difference, AT
however, shows a systematic and fairly large change from undisturbed to
disturbed conditions. Mean sea-air temperature differences in excess of 1C
occur during the disturbed modes, while under the strong trade regime AT
is less than 0.3C. Sea surface temperatures were observed to vary little
between periods, although slightly lower temperatures were noted during the
strong trade wind regime as opposed to the weak regime. The dominant control
upon AT is produced by variations in air temperatures. Lower air temperatures
during the disturbed modes are ascribed to reduced insolation due to cloud
cover, direct cooling by rain showers, as well as the descent from aloft of
cool air associated with the condensation-precipitation cycle. These processes

114

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09°0 AT eam

are elearly not advective in origin, but are associated with the dynamics

of the disturbance involved. While these differences between the undisturbed
regime and the disturbed mode are noted, Figures 3 and 4 show that, with

the exception of AT in the disturbed case, there is little variation within
each mode or regime. The greater dependence of Ts - Ta upon wind speed in
the disturbed mode is related to the fact that much lower air temperatures
are observed at lower wind speeds under these conditions, as opposed to the
undisturbed regime. This leads to the reverse effect in Aq which is higher
at low wind speeds for the undisturbed case. The scalar average wind speed
is similar within similar intensity categories, i.e., the weak trade regime
speed is similar to the weakly disturbed mode wind speed. While it follows
from the above that, given similar 4q and u , similar Q, should be
obtained, it is also illustrated in Table 2 that the stability changes.

Under the strong trade regime the stability, as indicated by Rp, is on the
average, close to neutral. Maximum instability is reached in the weakly
disturbed mode, returning to weakly unstable in the strongly disturbed mode.

When a drag coefficient with a dependence upon stability is used, then these
changes in stability are reflected in changes in Cp. The changes in Cp

produce an increase in latent heat transfer during disturbed conditions even
though the wind speeds and specific humidity differences are similar to those

of undisturbed conditions.

Since there is a fairly systematic increase in AT there is a cor-
responding increase in sensible heat transfer during disturbed conditions.
Changes in the Bowen ratio show that this increase is greatest with respect
to latent heat transfer in cases categorized under the weakly disturbed mode.
This suggests that sensible heat transfer may play an important role in both
the formative stages of a disturbance and in the peripheral regions of an
organized disturbance.

The maximum exchange of total energy, Q , takes place in the disturbed
mode. The largest exchanges take place when conditions are closest to neutral
stratification and, in consequence, closest to conditions which are assumed
in the development of the exchange equations. Almost all of the relatively
large amount of total energy transferred from the ocean to the atmosphere
during undisturbed conditions is in latent form. It has been clearly
established by other workers such as Riehl (1954, p. 56) and confirmed in
this region by La Seur and Garstang (1964a), that the greater proportion
( > 50 per cent)of tropical precipitation falls in organized synoptic
disturbances. The proporation over the oceans may, in fact, be far larger
than this figure. Hence, before a significant proporation of the latent
heat accumulated in the trepics can be made available to the atmosphere, it
must be advected into organized synoptic disturbances. Condensation and
precipitation processes will release part of this latent heat directly into
the disturbance. This available energy may then be used both to fuel the
disturbance itself, as well as increase the potential energy in the upper
levels of the tropical and equatorial atmosphere. Therefore, the synoptic
disturbances, not only represent a localized maximum of energy flux, but
are also regions of horizontal convergence and vertical transport of the
energy supplied to the atmosphere over large regions of the tropics. The
extent to which individual synoptic disturbances contribute to the energy
budget of the atmosphere is examined below.

118

Jog pue sepow peqingstp ey} [Te tog peeds pura jo uoTqounjz @ se LV

JOJ pues Sepoul poqingstp e494 [Te Joy peeds puyzm jo uoTyouny ev se

10.0

90 95

“——~—=—_ UNDISTURBED

—————=—8 OiSTURBED

40 45 80 $5 60 65 70 75 @0 45

20 25 30 35

ui MOD OF AOS se NS Sen Sal oe

= a G
(8178) ° —“b 3oNauadsIO.NNV3~I

WIND SPEED (m/sec)

*SoUTsZa1l paeqingstpun 3244 [Te

*SouTZe1 peqingstpun oy [Te
by :(eaoqy) € aandty

60 65 90 95 i00

UNDISTURBED
50 55 60 65 70

s—= DISTURBED

40 45

19 20 25 30 35

Lan =p Momma as em eam
=P em st A OM 1). 4'O. 30 O Om— oO.

(De) 91 -*) 30N3U3III0 NVON

02
01

>(MOTe8d) 7 e1naTT

WIND SPEED (m/sec)

alaly)

The temperature and specific humidity distributions shown in Figures
3 and 4 suggest that the time sequences presented up to this point may,
given the wind field, be extended into space sequences. Table 3 shows
constant values for six conditions. The spatial variation will, as pointed
out above, depend upon adequate representation of the wind field. This was
done by compiling composite maps for each of the disturbances considered.
Surface streamline analysis, utilizing both reconnaissance and research
aircraft reports and satellite pictures (in 1963), was carried out for each
storm at intervals of 12 hours. Each 12-hourly chart consisted of a composite
of at least 6-hourly reports. Three-hourly reports were added when available.
In each of the disturbances the CRAWFORD was located well within the circula-
tion of the system. The time section and detailed surface reports at the
ship were incorporated into the analysis. A series of 12-hourly streamline
analyses was then obtained for each vortex. A period, which varied between
12 hours and 102 hours, in which the storm approximated a steady state at
the surface was selected. A composite of each storm was then constructed.
This was done by locating all observations on the individual streamline
analyses with respect to the center and neutral point. The observations
were then transferred and relocated on a composite map. An isogon-isotach
analysis was carried out for each composite map and the resulting streamline
isotach analysis used to perform a component analysis, from which the low
level field of divergence was computed. This was done for three cases. Too
few observations prevented detailed kinematic analysis of the fourth case,
but the streamline field could be delineated and the relative motion of the
vortex shown with respect to the stationary ship. Since this happened to be
the only vortex of which the center moved past the ship on the equatorial
side, it is the only example where flux measurements could be made through
the strong wind regions of the disturbance.

Two examples of the composite streamline Pields and the associated
divergence and weather are presented in Figures 5 and 9. The values of Aq
and AT given in Table 3 are showm in Figures 6 and 10, each region being
defined by one or both of two criteria:

1. The speed field (limits given in Table 3);

2. The division between disturbed and undisturbed conditions given
by the line of zero divergence and by the distribution of the weather.

The energy transports associated with the first example are shown in
Figures 7 and 8. The dependence upon wind speed controls the major features
of the distribution of Q,, maximum values occurring within the speed maxima
around the center and to the north of the center, while minimm values of
Qe are associated with the speed minima around the cols and center. Within
the regions of maximum latent heat transport, values of Q. are associated
with the epeed minima around the cols and center. Within the regions of
maximum latent heat transport, values of Q, exceed 600 cal em~day-1 over
fairly large areas. These large transports are concentrated within the
regions usually associated with maximam precipitation. Five centimeters of

120

TABLE 3

AIR-SEA PROPERTY DIFFERENCES BASED UPON TABLE 2 AND FIGS. 3
AND 4, USED TO COMPUTE ENERGY TRANSFERS FOR THE COMPOSITE

DISTURBANCES
a = -1
AN CGY Ng, Cs Keli
Strong trade 0.20 560
Jol im Seer
Moderate trade 0.30 Seid)
So0=7o/ im SEe7
weak trade 0.40 Bo 8
<€ 5.0 m SEQ”
Weakly disturbed A iL LO Hoe
< 3,5 im SeeC~
Weakly disturbed B 0.80 B50!
3.5-5.25 m sect
Moderately and strongly 0.65 Sie

disturbed

= S525 im sec7t

wef.

Streamlines (solid), divergence (dotted in units of
1072sec7!) and weather based upon observations

composited over the period from O000 GMT on 15 August
to 0600 GMT on 19 August 1957.

122

Figure 6.

° 0° ° 2 3° ra

Isotachs (solid) and regions of constant Aq and
AT . (dotted) based upon (a) the delineation of
the distrubed region of the storm using the zero
line of divergence and distribution of weather

in Figure 5 as a guide; (b) the classification in
Table 3. AT is the upper and Aq is the lower
figure in each region.

123

1e o 0 es 3

ad
a

( eat nt) =

1 = 4 L 3s A 7 1

Figure 7. Latent heat transfer in units of cal em™@
day” - based upon Figures 5 and 6. The central
values represent the highest and lowest value
computed in each closed region.

12h

day"! based upon Figures 5 and 6. Notation as
for Figure 7.

Figure 8. Sensible heat transfer in units of cal cm”

125

Q=-4Q,» 0-2 i=
Caw

Figure 9. Streamlines seen divergence (dotted in
units of 10° sec and weather based upon
observations nett Ra over the period 0600
GMT on 21 August to 1200 GMT on 21 August 1963.

126

Figure 10.

Lod +? LD 2 we oI an 3 o 8° e@ uw e

Isotachs (solid) and regions of constant Aq

and AT (dotted) based upon (a) the delineation
of the distrubed region of the storm using the
zero line of divergence and distribution of
weather in Figure 9 as a guide; (b) the classifi-
cation in Table 3. AT is the upper and Aq is
the lower figure in each region.

Figure 1l.

eee

Latent heat transfer in units of cal em7< day~1
based upon Figures 9 and 10. The central values

represent the highest and lowest value computed
in each closed region.

20.

127

128

Figure 12. Sensjble heat transfer in units of cal om
day ~ based upon Figures 9 and 10. Notation
as for Figure ll.

129

rain in such a region in 1 day would represent a release of about 3000 cal
em~“day~_. While the greater fraction of water vapor represented in such
an energy conversion mst be due to horizontal advection, a significant
amount could be accounted for by the transfer processes taking place within
such a region. The distribution of sensible heat transport is similar to
that of the latent heat transport, with the exception that strong gradients
coincide with the boundary between disturbed and undisturbed states. Since
this is the boundary between the regions of maximum and minimum cloudiness
and weather, maximum values of Qs are found within the disturbed region.
Values of Qg in excess of 36 cal cm-“day~- are observed over a wide crescent
around the center of the disturbance. This means that maximum convective
instability should be concentrated within this region. Total energy Q is,
therefore, at a maximum within this region, exceeding 650 cal em-2day-1 in
the northern quadrant of the storm. The distributions of Qe and Qs for the
second example are shown in Figures 11 and 12. The main features are un-
changed, but there is considerable increase in the size of the transports.
Latent heat transport reaches a_maximum of 1446 cal em=“day-1 | and sensible
heat transport, 78 cal em” “day” - ‘These values can be compared with the
maximum values computed by Petterssen, et al., (1962) from typical winter
cases of cyclones in the western North Atlantic. Maximum values of 1440 cal
em-2day-1 for latent heat flux and 720 cal cm-@day-1 for sensible heat flux
were obtained. Pyke (1964) computed valugs for a cyclone in the Gulf of
Alaska and obtained values of 350 cal em~“day~! for Q. and 216 cal cm-“day-1
Por - Malkus and Riehl (i958), assuming equations essentially the same

as (3 and (2), computed latent and sensible heat fluxes for a moderate
hurricane within 90 km of the center of 2420 cal cm™ day~t and 720 cal
em7“day~ - Consistent with the above, they assumed small changes in dg - qa,
a decrease of 5.2 g ket in the region 90 to 70 km to 3.5 ¢ kg-l 50 to 30 km
from the center. Thus the increase in latent heat from the values observed
in the above disturbances and in the trade wind regions is due to an increase
in wind speed. They assumed a constant sea-air temperature difference of
_2.0C. This they ascribe to adiabatic expansion during horizontal motion
toward lower pressure rather than the cooling mechanisms called upon in the
above disturbances. This large AT , together with high wind speeds, gives
a large sensible heat transfer. In comparison with this value, ges highest
hourly mean value of Qs observed on the CRAWFORD was 201 cal cm” day™-.

The values utilized by Malkus and Riehl for the moderate hurricane are,
therefore, quite consistent with those values computed from the composite
storm data. In turn, these are consistent with values of latent heat trans-
fer computed for higher latitudes. The values of sensible heat transfer
appear, at the most, to be 40 percent of the values observed at higher
latitudes where pronounced sea-air temperature differences are observed.

Figure 13 shows the observed values of latent and sensible heat transfer
every 2 hours through a composite streamline field of an equatorial vortex
SEGPERNE SSE) during the 1963 cruise. Latent heat flux ranges from 252 to 938
cal em7“day~1, while sensible heat flux ranges from -2 to 201 cal cm-@ day71,
Regions of maximum and minimum transfer coincide with the distribution noted

in the composite models presented in Figures 5 through 12.

130

Figure 13.

e e o

Streamlines based upon observations composited over

the period 0000 GMT on 25 August to OOOO GMT on

27 August 1963.

Observations made on the CRAWFORD
are plotted every 2 hours with the corresponding
values of latent and sensible heat in cal cm “day”

on the upper right hand side of the plotting model.

131

The integrated transport of latent and sensible heat was computed for
all disturbances and mean values for the whole area are presented in Table h.
These mean values are compared with values which would have been obtained had
the whole region been covered by a uniform trade wind, maintaining an average
speed within the limits specified, for periods of 1 to 4 days. Climatic
charts for the Atlantic from 15°S to 15°N from September to November
(Mac Donald, 1938) show that average wind speeds are everywhere less than
6.7 m sec71, and over most of the region less than 5.0m sec7+. From Table 4
the average value of Q. would be between 180 and 380 cal em~“day-1 and of
Qs between 5.4 and 10.5 cal cm-2 day-+. The occurrence of a moderate synoptic
disturbance of dimensions similar to those considered above would greatly
alter the distribution of energy input within these latitudes. Only during
the months of November through May. do average winds in excess of 7.2 m sec”
cover a significant portion of the region 15°N to 15°S. At this time of the
year within the moderate to strong trade wind regions, the latent heat flux
would approach that associated with moderate to strong disturbances. For
the remainder of the year, i.e., June through September, the net input of
energy from the surface of the ocean in this region will depend significantly
upon the frequency and distribution of such synoptic scale disturbances.

Detailed analysis of the tropical Atlantic for the International
Geophysical Year of 1958 is being carried out by Dean ( Dean and La Seur,
work in progress in Dept. Meteor., Florida State Univ.). The frequency of
equatorial vortices for each month in 1958 has been obtained. Preliminary
results suggest a frequency of 10 to 15 vortices per month during August,
September, and October. As shown in Figure 14, development generally takes
place off the West African coast in the vicinity of 5°n to 107, with some
vortices apparently emerging from the continent itself. As the equatorial
trough region migrates equatorward, the number of vortices diminishes and
they can be tracked into the Atlantic for limited distances only. During
the months of January and February, only a few weak vortices appear in the
eastern equatorial Atlantic, none of which migrates into the central
Atlantic. Phenomena in the mid- and upper-troposphere (such as the tropical
easterly jet) appear to play an important role in this seasonal fluctuation
in the frequency of these equatorial systems. Figure 14 shows the tracks of
11 equatorial or tropical cyclones. At least nine of these systems followed
a track confined to an extremely narrow belt. As shown by Figure 15, the
mean map reflects this concentration as an asymptote in the streamline field.

For the region near the asymptote, a mean number of 10 vortices per
month can be accepted for the wet season months (July through October), with
1 to 4 per month during the dry season (January through May). If, in the
mean, the disturbance covered 15 x 15 degrees of latitude and moved at a
meen rate of 300 nautical miles within 24 hours, or 5 degrees of latitude
per day, then within a belt 15 degrees wide across the tropical Atlantic
centered on the mean asymptote during any wet season month, the mean
sensible and latent heat flux would correspond to the values computed for
the composite storms. The choice of a belt 15 degrees wide is suggested
not only by the mean size of the disturbances examined, but is also reflected

132

TABLE 4

COMPARISON BETWEEN LATENT AND SENSIBLE HEAT TRANSFER UNDER
CONSTANT TRADE WIND CONDITIONS AND SYNOPTIC SCALE DISTURBAN _
CES OVER AREAS OF 10 BY 10 AND 20 BY 20 DEGREES OF LATITUDE

Average Energy

Fluxes

in the

Trade Wind
Region

Average Energy Fluxes
in Disturbed
Conditions

Disturbance Disturbances

On a.

Weak trade

Q.<180
U<2.5 im See

Weak to
moderate
trade
2.5<u<5.0

180<Q,<314

Weak
disturbance

Moderate
trade
5 O<U<7 57

314<Q,<450

Moderate
disturbance
< tropical
storm

Strong trade Qe 7450
WE > 7/

Strong disturbance

< hurricane

Q.<55%

5.4<Q.<9.4

Q,heOa<siIa 8

I ie e@ayel ICICI
averaged ‘averaged
over , over
100°lat 400° lat?
418 ES
435 17
480 22

Figure 14.

55 oF 45 O 35 0 25 20°

Tracks of equatorial cyclones for the month of
August 1958. The cyclones have been identified
in the surface streamline analysis and tracked
until the circulation could no longer be iden-
tified. Tracks of systems originating in July
and persisting into August have not been in-
cluded. (After Dean, Dept. Meteor., Florida
State Univ.)

133

ss° $o° as ae 3° 30° 2s° 20° 1s” to°

Mean streamlines (solid) and mean wind speeds,
irrespective of direction (dashed) for the tropi-
cal Atlantic based upon all ship reports within
5 degree latitude-longitude squares for August
1958. (After Dean, Dept. Meteor., Florida

State Univ.)

30°

25°

20°

s°

135

Mean streamlines (solid) and mean wind speeds,
irrespective of direction (dashed) for the tropi-
eal Atlantic based upon all ship reports within

5 degree latitude-longitude squares for February

1958. (After Dean, Sept. Meteor., Florida State
Univ.)

136

by the charts of wind constancy (e.g., MacDonald, 1938). For the months of
July through October a belt 15 degrees wide with wind constancies of

80 percent or less exist about the equatorial trough zone. A constancy of

80 percent signifies that during 80 percent of the time steadiness is better
than 90 percent. Therefore, the 80 percent constancy boundary roughly corres-

ponds to the division between disturbed and undisturbed conditions used above.
Outside this belt other synoptic systems would affect the distribution of

energy flux, but within the stronger and more persistent trade wind belts

such fluctuations would have less effect upon the net transport than is the
case in the vicinity of the equatorial trough.

The month of August was chosen as representative of the wet season
and the month of February as representative of the dry season. Computations
for the tropical Atlantic have, in each case, been based upon the mean wind
speed values depicted in Figures 15 and 16. On the August map the 15-degree
wide region under the influence of disturbances has been centered on the
position of the mean asymptote. Latitudinal means were obtained for Qe and
Qs at 1-degree intervals from the values computed for the composite storms.
These values were applied to 2h of the days of August, the remaining 7 days
represent undisturbed conditions. The region between +7.5 degrees of the
asymptote and about 30°N and 20°S was then subdivided on the basis of mean
sea surface temperatures (MacDonald, 1938). Based upon these values, the
air-sea temperature difference was scaled according to observed values from
the 1957 and 1963 data. The difference between the saturated specific humidity
and the humidity at 6.0 m was assumed constant at 5.0 g kg-l. Computations
(as before) of Qe and Qs were then made for 5-degree squares. This procedure
of computing Qe and Qs outside the equatorial trough region in August was
applied to the whole area in February. Transport during the transition
months between wet and dry seasons will depend upon the occurrence and
frequency of synoptic disturbances. Thus, an annual mean map of Qe and Qs
has been constructed from an average of the August and February maps. These
six maps are presented in Figures 17 through 22. The mean value of Qe for
the month of August shown in Figure 17 shows similarities with the mean annual
map of Budyko (1956) only near 30°N. On Figure 17 a well defined maximmm now ~
appears to the north of the asymptote as a result of synoptic scale systems.
A secondary maximum appears in the southwestern tropical Atlantic in associa-
tion with the maximum in the winter trade of that hemisphere. The overall
results indicate a pronounced increase in latent heat transport within the
equatorial and southern tropical Atlantic.

The mean sensible heat transport for August show in Figure 18
indicates, in general, lower values than reported by Budyko (1956) but higher
values than Jacobs (1951). Within the equatorial trough regions values are
close to those shown by Budyko. This is due to the added transport associated
with synoptic systems. Outside this region the values reported in Figure 18
are lower than those of Budyko. This is in large part due to the exclusion
of diurnal effects, a factor not considered by Budyko.

During February the mean transport of latent heat shown in Figure 19 |
corresponds in the southern hemisphere to that obtained by Budyko (1956) on

Mean values of latent heat transport in cal cm.
day-1 for the tropical Atlantic for the month of
August, computed on the basis of methods outlined
in the text.

Ushi

138

Figure 18. Mean_values of sensible heat transport in cal cm
day"! for the tropical Atlantic for the month of
August, computed on the basis of methods outlined
in the text.

139

20°

10°

o*/}-— o°

10°

20°} -

Figure 19. Mean_values of latent heat transport in cal an
day~1 for the tropical Atlantic for the month of
February, computed on the basis of methods out-
lined in the text.

140

80° 70° 60°

By
Ge
10°) a —— a =
\
NS
‘I
20°}; ——— =
i
ao Oe 7o° 60°
Figure 20. Mean, values of sensible heat transport in: cal ous

day ~ for the tropical Atlantic for the month of
February, computed on the basis of methods out-
lined in the text.

141

Figure 21. Mean Baed. values of latent heat transport in
cal cm~“day"! for the tropical Atlantic based
upon Figures 17 and 19.

142

Figure 22. Mean annual values of sensible heat transport in
cal cm7“day~! for the tropical Atlantic based
upon Figures 18 and 20.

: : 143
an annual basis in the northern tropical Atlantic. A maximum associated

with the trade-wind maximum of the winter hemisphere appears again.

A marked decrease in the transport of sensible heat during February is
shown in Figure 20. This is not only due to the inclusion of diurnal effects
put also to the fact that fair weather is a mean condition over most of the
region during this month.

Figures 21 and 22 show the mean annual distribution of Q, and Q, for
the tropical Atlantic as calculated in this paper. Im the case of latent
heat the most significant features remain the maximum in the equatorial
trough region associated with synoptic scale disturbances, and the higher
values shown over much of the southern tropical Atlantic. The integrated
effect still produces a maximum within the eastern and central equatorial
trough region.

The inclusion of both the synoptic and diurnal effects on the transport
of sensible and latent heat produces pronounced variations in both the monthly
means and the annual distributions. The inclusion of the drag coefficient as
a function of both wind speed and stability results in an increase in the
transport of heat, particularly in the regions of high mean wind speeds.
Clearly, these effects will apply in all tropical ocean areas. In the central
and western Pacific, where the mean annual frequency of synoptic disturbances
is higher than in the Atlantic, the effect will be even more pronounced than
shown to be the case in the Atlantic.

IV. CONCLUSIONS

The strong dependence of latent and sensible heat transport upon
synoptic scale disturbances implies that the role of the equatorial trough
region in the heat balance of the atmosphere may be even more important than
has heretofore been suspected. WNot only do the disturbances in this region
perform the function of horizontal and vertical advection of vast quantities
of energy generated within the trade wind regions, but a large increase in
energy transfer occurs within the region of the disturbance. Through the
condensation-precipitation cycle, the disturbance then provides a means of
releasing this energy. In the mean, the concentration, vertical and horizontal
transports of energy in this region prescribe one of the fundamental con-
straints on the heat balance of the atmosphere.

While it is considered that the magnitudes of the transports derived in
the text may be more accurate than those previously published, it is important
to note that changes in the coefficients used would not change the direction
of the transport, the sense of the synoptic variations or the fact that
synoptic systems produce large amounts of latent and sensible heat at the
surface. Verification of the absolute magnitude is indeed desirable but there
seems little doubt that a region of maximum transport will occur just poleward
of the equatorial trough and that this region will have maximum values in
excess of those appearing on current maps (e.g., Budyko (1956)). These con-
clusions cast some doubt upon the relative magnitude of the various terms in

14h

the presently accepted heat budget of the ocean and the atmosphere. In
particular, it may suggest a revision of the current estimate of the radiation
balance and oceanic heat flux (Budyko, 1956). However, before such a
revision could be made, the uncertainties that have entered into the con-
putations made in this paper would have to be removed.

Bellamy, J.C.

Budyko, M. I.

Deacon, E. L.; P. A. Sheppard;
and E. K. Webb

3 and E. K. Webb

Garstang, M.

Jacobs, W. C.

La Seur, N.E. and M. Garstang

145

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1956:

1962:

1958:

1961:

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1964:

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The heat balance of the earth's
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146

MacDonald, W. F.

Maikus, J. S.

and H. Riehl

Monin, A. S. and A. M. Obukhov

Petterssen, S.; D. Le Bradbury and
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Riehl, H.

and J. S. Malkus

3 T.-C. Yeh; J. S. Malkus;
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147

INTENSITY OF HURRICANES IN RELATION TO SEA
SURFACE ENERGY EXCHANGE

. Irving Perlroth
National Oceanographic Data Center, Washington, D.C.

148

ABSTRACT

The following study, based on the apparent relationship existing
between variations of central pressure of a hurricane and the sea-surface
temperature pattern, indicates that hurricane intensity is affected by the
eddy flux of energy from sea to air. The effect is investigated by con-
structing synoptic composite sea-surface temperature and related energy
exchange charts, employing Jacobs' equations for determining the energy
removed from the sea.

149

INTRODUCTION

The relationship between sea-surface temperatures and the intensity
of hurricanes is becoming more evident with the increasing comprehension of
the oceanographic environment. Recent studies by Fisher (1957), Perlroth
(1962), and Tisdale and Clapp (1963) indicate that energy exchange between sea
and air is one of the major factors governing hurricane structure. Namias
(1962) found similar evidence in studying the behavior of typhoon Freda in
September 1962. For the following study on the effects of energy exchange
upon the behavior of hurricanes, the author has selected hurricane Ginny
(October 1963) as the main example to serve his purpose.

To fully understand the interaction between the sea and the atmosphere,
we need a detailed analysis of the sea-surface temperature field. Due to
the inherent difficulties in obtaining sea-surface temperatures in the im-
mediate vicinity of hurricanes, a composite data anlysis was performed im-
ediately before the storm's passage.

Previous studies by Gibson (1962) and Perlroth (1962) have shown that
construction of 10-day composite sea-surface temperature charts with definable
temperature patterns is possible. It is believed that the extent of order
and stability of 10-day composite sea-surface temperature patterns permit
such charts to be representative of the steady-state structure of the
temperature field for any given day during the indicated period.

In this study attempts. are made to relate the fluctuation of hurricane
intensity with the sea-surface temperature pattern. No conclusive evidence
is present that a hurricane follows a track which lies over the warmest water
(Fisher, 1957); however, hurricane Ginny (October 1963) appeared to traverse
the core of the Gulf Stream for a long time. It is believed that under
unusual atmospheric conditions (devoid of any major steering currents), the
paths of hurricanes may be influenced by areas of maximum energy exchange.
The cycloidal character of many hurricane tracks might be attributed to the
influence of pronounced sea-surface temperature patterns.

It is also believed that energy obtained from the sea is not the sole
controlling factor governing hurricane intensity and fluctuation of central
pressure; however, if a hurricane remains tropical in its character and is
not extensively influenced by any cold fronts or extratropical troughs, there
appears to be substantial evidence of the effect of energy exchange on
hurricane structure.

150
HEAT-EXCHANGE COMPUTATIONS

Data for the analyses of Figures 1 and 2 were obtained from ships,
located in the western Atlantic, which transmitted radio teletype reports
of synoptic weather observations to the U. S. Weather Bureau. The basic data
used were air and sea temperatures, dew point, and wind velocity.

Jacobs' (1942) final equations for the energy removed from the sea are
as follows:

a, = LEdot (Ge. o Ss) te & cal/em@ day

Q. = 0.01(-t,, - t,) Q@ @ cal/em™ day

Cc —— eee
ew ~ a)

Q@, = 145.4 (ey - e,) 0.01 (t,, =u, } He,

where

Q, = energy used for evaporation

Q. = sensible heat exchanged between sea and atmosphere
through convection

Q, = Sum of Qe and Qc , representing total energy ex-
change between sea and atmosphere

€, = vapor pressure at height a, in inches
e,, = vapor pressure at sea surface, in inches
W, = Wind speed at height a, in knots

‘Ge = sea surface temperature, in degrees F

= air temperature at height a, in degrees F

Since in many previous studies computations for total eddy transfer of
heat have been averaged over large ocean areas (5- and 10-degree squares),
only generalized estimates can be made of the physical processes that may
exist; the resulting relationships therefore are incomplete. As pointed out
by Jacobs, the constants in the above equations are intended to apply only
to the marine climatic data which were used by him in his computations of
seasonal values over the oceans. Nevertheless, his constants have been
used subsequently in the more-or-less synoptic sense by a number of in-
vestigators; the results have almost invariably provided the proper order of
magnitude and have satisfied continuity, except when the constants were used
under conditions of extreme atmospheric instability. It should be pointed
out, however, that the validity of the constants has not been established

Bl /°82
7 OTT
Vw. 1179

84
83,9 at F

Figure 1. Sea Surface Temperature Pattern
11-19 October 1963

152

Fe |220
1270 lOo7C

aig

870 | (|
@ A 7

°960 ( 900
1020

°720

Figure 2. Heat Exchange Pattern 11-19 October 1963
(Values in hundredths g cal/em® day)

153

when the computations involved the excessively high wind speeds found in
the vicinity of hurricane centers. Figure 2 represents an attempt to
construct a composite heat-exchange chart for the 10-day period just before
the passage of hurricane Ginny. The pattern shows only the existing heat
exchange, unaffected by the passage of the hurricane. Dew points, air-sea
temperature differences, and wind velocity were averaged for this 10-day
period by l-degree squares. Sea-surface temperatures and air temperatures
for each synoptic observation were used in the calculation of the total
heat exchange in each l-degree square. By employing this method, a composite
energy exchange chart, based on a corresponding composite sea-surface
temperature analysis, was constructed (Figure 1). Patterns of warm and cool
water masses and of maximum and minimum energy exchange areas are shown in
Figures 1 and 2.

Extensive calculation errors which may occur in the preparation of
total heat-exchange charts can be minimized by using data from composite sea-
surface temperature charts. The most obvious errors in determining sea-surface
temperatures can be eliminated by mass data coverage (Figure 1). Consequently,
Jacobs' equations can be used in computing individual synoptic data, and
reliable results can be obtained.

OCEANOGRAPHY AND ANALYSIS

To construct the 10-day composite chart shown in Figure 1, sea-surface
temperature data for October 11-19, 1963, were used. An analysis of these
data was performed; doubtful data were circled and not used in the computa-
tions. The chart shows that the temperature patterns for certain ocean areas
maintain a high degree of persistence. For shorter time periods (10 days or
less), these patterns appear to be conservative; they are, however, extremely
complex. Berson (1962) endorses the existence of quasi-meridionally oriented
bands of thermally differentiated water, with significantly varying layer
depth, having dimensions of 20 to 60 km in width and 500 km or more in length.
The surprisingly stable correlations with current velocity suggest long-term
persistence of these bands. This persistence is in agreement with strong
indications of quasi-geostrophic balance prevailing in the long-term seasonal
averages of the current component along the bands. It is thus quite possible
that these standing oceanic eddies, on a mesoscale, form an integral part of
the mechanism for secular-scale meridional heat transport in the oceans.
According to Laevastu (1963), pronounced patterns of heat-exchange components
are relatively persistent from day to day, with slight changes in position
and intensity.

Figure 1 shows the axis of the Gulf Stream and the adjacent cold water
of the continental shelf. Horizontal surface temperature gradients of 8 to
10°F exist within the few miles between these two water masses. The alternat-
ing warm and cool bands of water are quite notable; however, their surface
temperature gradients are less than those along the Gulf Stream.

The total-heat-exchange pattern shown in Figure 2 is very similar to

154

the sea-surface temperature pattern. Computations of total ener flux

(sea to air) in the core of the Gulf Stream exceed 1600 g cal/em* day; yet
only a few miles seaward, approximately half of this value was obtained. The
maximum heat-exchange areas coincide with bands of warmer water, and minimum
heat-exchange areas coincide with bands of cooler water.

The absolute values of sea-surface temperature and heat exchange undergo
a change upon the passage of a hurricane; however, the relationship between
temperature and heat exchange remains the same during the passage of a hur-
ricane. In the Gulf Stream, near the centes of hurricane Ginny, calculations
of total heat exchange exceed 3000 g cal/cm< day.

Bathythermogram studies of the vertical temperature structure of the
ocean in the numerous water bands have indicated interesting relationships.
The warmer water bands, particularly the core of the Gulf Stream, are generally
isothermal to 150-250 feet during this time of year (October); the cooler
water bands ‘show a mixed layer of less than 50 feet. As the hurricane passes
over the cooler water, a notable decrease of the surface temperature can be
expected, while the deeper layers of isothermal water are less affected. It
could be concluded, therefore, that the surface temperature gradients do not
dissipate as is show by Figure 6. The sea-surface temperature and hurricane
pressure curves indicate a unique relationship.

HURRICANE GINNY

Hurricane Ginny formed in the western Atlantic, in the vicinity of the
southeastern Bahamas, on October 16, 1963. In her formative stages, this
storm appeared to be extratropical in nature. Figures 3 and 4 represent the
surface weather chart and 500-mb chart for 12 UT on October 19. The available
data at 500 mb indicate the presence of a deep polar trough along the south-
east coast of the United States, with an apparent cold pool of air in its
axis. However, transformation of the cold core to a warm-core low had ap-
parently occurred by October 22 when the reconnaissance aircraft found that
the storm had acquired tropical characteristics.

For study purposes, hurricane Ginny might be classified as an "ideal"
storm, primarily because of its erratic motion, slow movement, and the
coincidence of its track with the axis of the Gulf Stream. The "Yankee
Storm" in November 1935 followed a similar track. It developed in the
vicinity of Bermuda, moved westward to the vicinity of Cape Hatteras, and
from there curved southwestward to the south coast of Florida.

In this evaluation, the track of hurricane Ginny during October 22-28
is being studied (Figure 5). Hurricane positions indicated in Figure 5 were
obtained from the U. S. Weather Bureau radio teletype summaries of reconnaissance
aircraft penetrating the eye of the storm. It is notable that a considerable
increase in the intensity of the storm occurred between October 24, 1400 UT,
and October 25, 0100 UT, when the eye of the hurricane entered the core of
the Gulf Stream. The increase in intensity was apparently due to the increase

155

45° Nae 7
. H
i YS
1
40°

Figure 3. Surface Weather Chart 12Z 19 October 1963

156

30°

85° 80° 75° 70° 65°

Figure 4. 500 mb Weather Chart 127 19 October 1963

157

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iveral

158

in sea-surface temperature and the resulting increase in total energy exchange;
the eye and wall cloud of the hurricane became better defined, and the central
pressure dropped 7 mb in 7 hours. The intensification continued as the storm's
track lay over the axis of the warm water stream.

Dunn and Miller (1960) indicate that in the absence of mid-tropospheric
ventilation, the vertical temperature structure of the hurricane core is
determined by the heat and moisture content of the subcloud air which, in turn,
is closely related to the temperatures of the underlying water surfaces. For
instance, hurricane Janet (1955) deepened over warm water, and hurricane
Carrie (1957) weakened over cool water; however, the rate of change in hur-
ricane intensity depends upon the length of time it takes a hurricane to
traverse a particular water mass.

Riehl (1963) points out that a relationship exists between the develop-
ment of hurricanes and the small anomalies of local temperatures. Perlroth
(1962) exemplifies this relationship between hurricane intensity and the air-
sea interaction in a comprehensive study of hurricane Esther (1961).

The sea-surface temperature and hurricane-pressure relationship is
illustrated in Figure 6. For this relationship to be coincident, the sea-
surface temperature pattern during the 10-day period must have remained con-
servative. As shown in Figure 5, when the hurricane eye and wall cloud entered
part of the cold shelf water on October 25, a marked decrease in intensity
occurred. The central pressure of 976 mb was observed at October 25, 1200 UT,
while by October 25, 1859 UT, .a reading of 985 mb was obtained. The storm
then turned sharply to the east-northeast and once more entered the axis of
the Gulf Stream. The central pressure of hurricane Ginny again began to
deepen and by October 26, 0359 UT, reached 983 mb. On October 26, the hur-
ricane appeared to be headed for Wilmington, N. C.. Once again, as the eye
passed over the cold shelf water, a notable decrease in intensity was observed.
Aircraft reconnaissance indicated that the wall cloud was becoming diffuse,
and a central pressure reading of 988 mb was observed. These observations
indicate that the storm structure responds spontaneously to the total energy
flux from sea to air in the vicinity of the hurricane eye.

On October 28, the steering of hurricane Ginny became influenced by a
deepening polar trough forming along the east coast, consequently causing a
rapid acceleration of the hurricane to the northeast. Because of the influence
of this trough, correlations of hurricane intensity and sea-surface tempera-
ture were not determined. The hurricane reached the maximum intensity east
of New England as it accelerated rapidly north-northeastward.

CONCLUSIONS

The results of this study imply that a relationship exists between
hurricane intensity and the total energy flux from sea to air. The evidence
produced in this and other studies brings out the significance of this
relationship.

159

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It is apparent that where sea-surface temperatures are relatively low
(weaker energy exchange), hurricanes are cut off from their heat source and
tend to weaken; where the sea-surface temperatures are relatively high
(maximum energy exchange), hurricanes tend to intensify. (Figure 6 shows the
spontaneous response of hurricane central pressure to the sea-surface
temperature field.)

It is interesting that the facts brought out in this study would aid
the hurricane forecaster in predicting variations in hurricane intensity.
Construction of 10-day composite sea-surface temperature charts for the east
coast, Gulf of Mexico, or any other areas influenced by the presence of
tropical storms is desirable. The construction of total heat-exchange charts
ean be performed on a mesoscale, to provide the forecaster with a better
understanding of potential areas of hurricane intensification or weakening
along its projected track. It must be realized however, that outside the
major shipping lanes (tropical Atlantic) lack in density of synoptic observa-
tions may prevent a detailed analysis.

Numerous hurricanes, for instance the one of September 1938, of
September 1944, and hurricane Hazel in 1954, have reached maximum intensity
in the northern latitudes. These hurricanes maintained or even increased
their intensity after moving over colder coastal waters or over land. These
effects may be attributed to polar trough intensification which affected
the motions of these storms and also provided new sources of energy for
extratropical development. (This study is intended to relate only the
fluctuations of the central pressure of hurricanes in relationship with the
sea-surface temperature pattern of storms that remained tropical in their
nature. )

Hurricane Ginny appeared to follow a track over the Gulf Stream; how-
ever, it is believed that this track coincided with the prevailing weak
steering current. There is no substantial evidence to prove that the paths
of hurricanes follow the warmer water bands; for instance, hurricane Esther
(1961) assumed a track at right angles to the alternating warm and cool
water masses.

Although only a relatively small number of hurricanes has been studied
for determining the relationship of heat exchange with hurricane pressure,
there is sufficient evidence that hurricane intensity is related to sea-
surface temperature patterns. This concept should be fully investigated.

Acknowledgement - I am indebted to Dr. W. C. Jacobs for constructive
comments and criticism in preparing this study for
publication and to the U. S. Weather Bureau, Washington,
D. C., National Airport, for use of their files of
synoptic weather maps.

Berson, F.A.

Dunn, G.E., and Miller,
B. I.

Fisher, E. L.

Gibson, B.

Jacobs, W. C.

Laevastu, T.

Namias, J.

Perlroth, I.

161

REFERENCES

1962:

1960:

1957:

1962:

19h2:

1963:

1962:

On the influence of variable large-
scale wind systems on the heat
balance in the active layer of the
ocean, Tech. Mem. 25, National
Meteorological Center, U. S.
Weather Bureau.

Atlantic Hurricanes, Louisiana
State University Press.

Hurricanes and the sea surface
temperature field, the exchange

of energy between the sea and the
atmosphere in relation to hurricane
behavior. Report 8, Parts 1 and

2, National Hurricane Research
Project.

The nature of the sea surface as
deduced from composite temperature
analyses, Deut. Hydrograph. Z.

On the energy exchange between sea
and atmosphere, J. Marine Res.,

5(1), 37-66.

Imergy exchange in the North
Pacific; its relations to weather
and its oceanographic consequences,
Oceanography Division, Hawaii
Institute of Geophysics, University
of Hawaii.

Large-scale air-sea interactions
over the North Pacific from summer
1962 through the subsequent winter,
J. Geophys. Res., 68(22), 6171-6186.

(A) Persistence of composite sea
surface temperature patterns,
Undersea Technology, 3(4), 16-22.
(B) Relationship of central pres-
sure of hurricane Esther (1961)
and the sea surface temperature
field, Tellus, 14(4), 403-408,1962.

162

Riehl, H. 1963: On the origin and possible modifi-
cation of hurricanes, Science,
141 (3585).
Tisdale, C. F., and 1963: Origin and paths of hurricanes
Clapp, P. F. and tropical storms related to

certain physical parameters at
the air-sea interface, J. Appl.
Meteorol., 2(3).

163

THE GULF OF MEXICO AFTER HILDA (PRELIMINARY RESULTS)

Dale F. Leipper
Department of Oceanography and Meteorology
Texas A&M University

‘
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165

ABSTRACT

Hurricane HILDA crossed the Gulf of Mexico in the period September 30,
to October 4, 1964, developing to a very severe hurricane in the central
Gulf. Sea temperature data available prior to the storm indicated what is
probably a typical late summer situation with some surface temperatures
running above 30 C. Beginning on October 5, a 7-day cruise was conducted
over the area where hurricane winds had been observed. Using the Bureau of
Commercial Fisheries vessel GUS III, four crossing of the hurricane path were
made, one where the maximum 150 mph winds were observed, one south of that
where the winds had first reached 120, one north where they had decreased to
120, and one in shallow water (40 fathoms), where prior data had been collected
by the U. S. Fish and Wildlife Service from their Galveston Biological
Laboratory. Bathythermograms were taken regularly to depths of 270 meters and
hydrographic casts to 125 meters. All four sections of observations indicated
similar patterns of upwelling. During the passage of the hurricane it appears
that sea surface temperatures over an area of some 70 by 220 miles decreased
by more than 5 C, and that a cyclonic ocean current system was established
around this area. The data collected on the GUS cruise appear to be the first
systematic oceanographic observations available in such a situation. Although
they do not permit a full description of the changes which occurred, they are
suitable to serve as a basis for a model from which, for example, total amounts
of heat lost to the atmosphere might roughly be estimated.

INTRODUCTION

Most of the other papers presented at this Conference have dealt with
the influences of the underlying sea upon the atmosphere. This one illustrates
an exchange in the opposite direction, a situation where intense atmos-
pheric phenomena brought about significant and observable changes in the under-
lying sea. Although changes of similar type probably are caused by certain
less extreme weather conditions, it is seldom that features as distinct as
those created by hurricane HILDA may be observed.

Figure 1 shows the path of HILDA across the Gulf of Mexico. It may be
noted that the most intense stage of the hurricane occurred when it was
centered 250 miles offshore in waters of over 1000 fms depth. Thus, the
effects of the hurricane upon the sea would quite likely be similar to those
resulting from similar storms in the larger ocean basins. Entering the Gulf
of Mexico with winds less than 80 mph, the hurricane intensified to the 150-
mile stage and the winds again decreased to less than 120 mph during the
passage across the Gulf. The width of the zone having winds of hurricane
velocity is indicated in Figure 1. The average propagation speed of the
storm may be obtained from recorded dates and times.

166

29° |

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Figure 1. Positions are those of the center of the eye

at dates and times shown. Wind speeds indicated
at center positions are as reported by the
Weather Bureau in the advisory transmissions.
The dashed lines show the distances from the
center to which hurricane speed winds were
thought to extend. Latitude (north) and
longitude (west) are indicated on the margins.

167

On September 30, when HILDA entered the Gulf, the conditions encountered
were apparently those typical of late summer, the Gulf not having been
disturbed by any previous hurricanes nor by any widespread and severe northers.
The surface waters over much of the Gulf were of relatively uniform temperature,
approximately 29 - 30°C, as indicated in Figure 2. The temperature depth
structure, based upon a few bathythermograms, collected prior to hurricane
HILDA and upon the limited available climatic data, consisted of a well mixed
layer from the surface to approximately 60 meters depth with a normal seasonal
thermocline beneath. In Figure 2, the path of the hurricane during its
development stage may be seen to have followed along approximately the center
of the initially high temperature zone.

CRUISE PLANS

As hurricane HILDA became a severe hurricane on October 2, efforts
were made to locate a research vessel which could be used to make a survey
in the hurricane area immediately after the passage of the storm. The
Galveston Biological Laboratory in Galveston, Texas, was able to provide the
90-foot shrimp boat GUS III, and a decision was made on October 2 to begin
the cruise on October 6. The Laboratory provided the crew of the vessel
together with scientific observers David Harrington and Stewart Law and
Captain Jim McMurrey. From Florida State University came Reed Armstrong, a
graduate student under Dr. Robert Stevenson Accompanying the author from
Texas A&M University was chief marine technician Kenneth S. Bottom.

At the outset, there was no fixed cruise plan. The first objective was
to retrace a line containing observations made immediately prior to hurricane
HILDA from the R/V ALAMINOS of Texas A&M University. The plan, as completed,
is shown in Figure 3. When the point H-9 was reached the boundary of the cold
water area was reached and sea temperature conditions appeared to be at almost
prehurricane HILDA values. It was decided to make a section from there
perpendicular to the hurricane path. This section was made so as to pass the
position of the anchored buoy NOMAD from which regular observations were being
collected. The next section to the north was chosen across the path at the
point of maximum hurricane intensity. In the final run from BT 57 to 63,
Figure 3, a line was repeated along which observations had been made prior to
the hurricane by the Bureau of Commercial Fisheries. Most of the observations
collected on the cruise of the GUS III were made at times 5 to 10 days after
passage of the storm at the same station locations.

Although there apparently had been no systematic observations previously
under similar circumstances, there are indications in the literature that
areas of low sea-surface temperature are often found in the wake of a hurricane,
(Fisher, 1958, and Jordan and Frank, 1964). Hidaka, 1955, reported similar
cold areas and developed a theory which indicated that there would be con-
siderable upwelling in the center of a hurricane and that a cyclonic current

168

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would be formed around’ the low pressure area. Stevenson (in press) reported
on shallow water ocean conditions associated with hurricane CARLA in 1961.

A MODEL

The observations collected on the GUS III cruise may be fitted into a
simple model based primarily upon the concept of wind-driven current. V. W.
Ekman, 1905, concluded that the effect of the wind upon the sea surface is
to set up a surface current 45° to the right of the wind. As depth in-
ereases, the direction of the flow turns to the right and the speed decreases.
When the ends of vectors representing velocity at different depths are
projected on a horizontal plane they form a spiral which has been called the
Ekman Spiral. The net transport of water in this layer of wind-driven current
can be shown to be 90 to the right of the wind direction in the Northern
Hemisphere.

Figure 4 is a schematic diagram of a hurricane similar to HILDA. As
may be observed, the winds at any given time create an Ekman net transport
outward from the center.

Since the hurricane is moving and since the actual transport direction
and speed vary with depth the full pattern is more complicated. In all events,
the passage of a hurricane causes the surface layers of the ocean to diverge
from the center. The surface waters which are moved aside must be replaced,
and the only source of replenishment is from below. Thus, cold water from
considerable depths appears at the surface in the hurricane eye.

Figure 5 shows schematically the change in isotherms in a section across
the hurricane path. It allows a comparison of the isotherms as they would
appear in the normal summer situation in the Gulf with those which would
appear after a hurricane passage.

Since the upwelled water arrives at the surface only after the winds
have blown for a considerable time, these waters would lose little heat to
the atmosphere through evaporation and conduction. The major loss of this
type would be from the warm mixed layer of surface water upon which the winds
acted directly and which was pushed aside in the Ekman drift.

THE DATA

With this model in mind, the data may now be considered. The surface
temperature pattern prior to the hurricane was shown in Figure 2. The
comparable pattern after the passage of the hurricane is given in Figure 6.
The darkest shaded areas indicate warm water, above 28. These areas appear.
to represent the original warm surface layer which has undergone the normal
seasonal cooling through the 12 to 14 days involved. The cold upwelled

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Schematic Diagram of Hurricane and Associated Net Water Transport

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waters appear along the path of the hurricane, 23° water appearing where
temperatures of 29° had been observed. The cold band along the Gulf Coast
in shallow water is probably not a part of this upwelling pattern, but is
in part due to the fact that the time interval between observations in this
area before and after HILDA was 24 days compared to a time interval of less
than one-half that for the other data. Also, northerly winds in October
probably caused considerable cooling and mixing in these shallow waters.

Figure 7 shows the pattern of temperature change between Figures 2 and
5. The climatic atlases indicate a normal seasonal change of less than one
and one-half degrees per month at this time. Thus, the change brought about
by hurricane HILDA appears to be some 5o in excess of this.

To illustrate the depth of isotherms along the section where winds had
been highest, 150 mph, Figure 8 was prepared. Since the path of the ship
was deflected from a straight line by currents, the section was drawn along
a straight line determined by the observations on the western end and by
projecting on to this line in the eastern extent the positions of stations
near it. The portion of the section based upon these projected stations is
indicated by dashed lines for the isotherms. Several features of the section
may be significant. The warm surface waters have been pushed to the left
and to the right. A strengthening of the thermocline is indicated by the
closer spacing of several of the isotherms separating these uniform warm
water bodies from the upwelled water. The isotherms seem to rise rapidly
on both sides of the hurricane and the coldest water at the surface appears
slightly west of the hurricane path. The warm surface layer is deeper at
the eastern end of the section.

Salinities were obtained to a depth of 125 meters and these, together
with temperature values, permit the computation of the density anomaly
sigma t. In Figure 9, lines of equal value of sigma t are plotted against
depth and distance. The shallowest depth of the more dense water is noted
at the hurricane path. Near the surface, as show by the constant values of
sigma t, is a layer of mixed water. Proceeding both to the east and the west
from the path of HILDA, the depth of this mixed layer increases to a distance
approximately 100 miles from the path.

The temperature structure across section C may be indicated in a
different manner by copying the temperature depth traces from bathythermograms.
In Figure 10, the BT's from section C, as well as those from the other sections
surveyed on the cruise of the GUS III, are show. Water warmer than 25° is
indicated by the crosshatched areas. All of this water is seen to lie away
from the path to the east and west. This water is well mixed, the mixing
apparently having been caused by the hurricane winds and by the heat loss to
the atmosphere which lowered the temperatures of the whole water layer from
29) = ROO We By o B8e.

The coldest water is that near the path of the hurricane. The BT's
in this sector do not show the surface isothermal layers characteristic of
surface cooling and mixing but rather have the rounded characteristic which

175

After Hilda

a-
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Sea Surface Temperature Decrease in °C Before Hilda To

Figure 7.

96°
30°
26°
26°

176

BT: NO.

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120

(METERS )

140

160

DEPTH

180

200

220

240

260

280

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—_

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Location of Section

Figure 8. Depth of Isotherms in Section Across Path after Hilda

177

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179

is typical of recently upwelled water. The other three sections show
characteristics similar to section C.

In proceeding from Station 41 to the eastward along C (see Figure 3),
the heading of the GUS III was maintained without change. However, later
determination of position indicated a strong drift to the north as indicated
at Stations 51 and 52. Also, in the vicinity of Stations 56 and 57 there
appeared to be a set to the west. These drifts indicated the existence of
an unusual current. Since salinity observations had been made to depths of
125 meters it was possible to compute the density distribution to these
depths. Further, since temperature observations reached 250 meters depth and
since there was a close correlation between temperature and salinity values,
the salinities could be inferred from the temperatures at depths between
125 and 250 meters and densities could be computed here. From the densities,
the dynamic height of the sea surface above the 250 meter reference level
could be determined. The topography is indicated in Figure 11. Associated
with this topography would be a relative geostrophic current of approximately
1 knot. This, then, is the current related to the distribution of mass
established during the hurricane passage by the wind drift current.

Since there were some bathythermograph observations made from the
ALAMINOS and from the GUS III prior to the passage of HILDA, it is possible
to represent the local change in temperature at some six positions. The
superimposed temperature structures at each of these positions before and
after the hurricane are shown in Figure 12. The three upper positions,
which may be located in Figure 3, are in shallow water and the observed
change is what one might expect from mixing along with some cooling.

At the three positions in deep water, all of which were along section A,
(indicated in the insert on Figure 8) there are several general features.
In general, the depth of the mixed layer before the hurricane was less than
that after, and the temperature of the mixed layer was higher than it was
after HILDA. It should be noted that all three of these observation stations
were located east of the path in a region to which the warm layer of surface
water would have been displaced from the center of the hurricane.

Considerable further study is needed before the data collected may be
fully interpreted. Present plans call for the preparation of one brief
article about the surface temperature change alone, another somewhat along
the lines of the present paper but organized for final publication, one on
the estimated energy redistribution (which requires considerable additional
research on conditions existing before the hurricane) and, finally, one on
the duration of the cold spot and the cyclonic circulation which was observed
on the GUS III cruise.

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BEFORE HILDA

OTHERS AFTER HILDA

200
Figure 12. BT Traces at Six Positions, Showing Temperature Structure
Before and After HILDA at Each Position. (See Fig. 3
for locations)

182

ACKNOWLEDGEMENTS

The work of the author was supported by the Geophysics Branch of the
Office of Naval Research through the Texas A&M Research Foundation. The
vessel GUS III together with her crew and two observers were provided by the
Galveston Biological Laboratory of the Bureau of Commercial Fisheries,

U. S. Department of the Interior. Technical assistance was provided by
Larry Brennan and secretarial work was done by Lydia Fenner.

183

REFERENCES

Fisher, E. L. 1958 Hurricanes and the Sea-Surface
Temperature Field. Journal of
Meteorology, 15, pp. 320-333

Jordan, C. L. and Frank, 1964 Project Report, On the Influence

Neil L. of Tropical Cyclones on the Sea
Surface Temperature Field, Florida
State University. NSF Grant GP-621.

Hidaka, Koji and Akiba, Yoshio 1955 Upwelling Induced by a Circular Wind

System. Records of Oceanographic
Works in Japan, Vol 2, No. 1,

Mareh 1955.
Stevenson, Robert E. and The Modification of Water Temperatures
Armstrong, Reed S. by Hurricane CARLA. Florida State

University. In press.

Ekman, V. W. 1905 On the Influence of the Earth's
Rotation on Ocean Currents. Arkiv

for Matematick, Astronomi och
Fysik 2 (iz); pp. 1-52.

185

EVIDENCE OF SURFACE COOLING DUE TO TYPHOONS

C. .L. Jordan
Florida State University

187

Evidence of marked cooling in the surface layers of the ocean. as
shown by the preceding paper by Leipper. has been noted in association with
some intense typhoons in the western Pacific. As discussed in a recent
report (Up cooling of this type is often clearly indicated by 15-day mean
charts but the cooling shown between successive 15-day periods seldom exceeds
5 F. ‘Individual ship traverses through areas over which intense typhoons
have passed have indicated cooling of the same magnitude shown by Leipper.

The sea surface temperature given in routine synoptic reports are
subject to quite large errors 2X, but the errors tend to be systematic for
an individual ship. Consequently. the temperature changes indicated along
a ship track are usually much more reliable than those deduced by combining
observations from several ships. Observations of this type have been used
to present two cases of unusually low sea surface temperature in the western
Pacific (Figures 1 and 2). The data for the ship which passed through the
area traversed by typhoon Wanda of 1956 ( Figure 1) were taken less than

ol-12
euuly 3! ve
sim ——_ 780s

—
128 130 132 134 136 138°E

Figure 1. Ship traverses prior to and following typhoon Wanda of 1956.
The positions of the storm center and its central pressure at OOOO GMT
are shown along the storm track. Temperature values (in F) are given
along the two ship traverses with dates and times shown below the individual
reports. The additional reports of sea surface temperature were made on
July 30 and 31.

188

fo)
36 hours following the passage of the storm. The two reports of 72 F are
10° colder than those reported by the same ship some 250-300 miles to the

northeast and up to 1,°F lower than values reported by ships 150-300 miles
to the north and northeast 1 to 2 days earlier. The ship traverse across

the track of typhoon Nina of 1953 (Figure 2). which occurred about 48 hours
after the storm, reported temperature as low as 74 F. These values are some

18 4 5
130 132 134 136 138 140°E

Figure 2. Same as Figure 1 except for typhoon Nina of 1953.

10° lower than the climatological average for the area and season and up to
1)°F lower than reports from a ship traversing the area 150-200 miles to the
north about a week before the typhoon passage.

The extent of cooling of the surface waters brought about by a tropical
cyclone is undoubtedly related to storm intensity and to the vertical
temperature distribution in the ocean. Typhoons Wanda and Nina were unusually
large and intense and there is little doubt that most tropical cyclones do
not result in cooling of the magnitude suggested by Figures 1 or 2 (or by
the previous paper by Leipper). However, ship reports following tropical
cyclones suggest that rather marked cooling occurs in many cases. Fisher [37
noted cold pools in the sea surface temperature field following several a
hurricanes in the 1953-1955 period and ship reports in the western Gulf of

189

Mexico indicate a significant decrease in surface temperatures following
hurricane Carla of 1961. The extent of cooling in this latter case can be
judged by the following sea surface temperature statistics for the area
26-29°N. 90-96 W. In the 4-day period prior to &he storm passage there were
1] reports of temperatures ranging from 85 to 88 F, with a mean value of
86.0°F- In the 5-day period following the storm. the reports in the same
area ranged from 78°to 82°F with a mean of 79.6°F.

In the report cited previously /1/, the conclusion was reached that
vertical mixing was the primary factor in the cooling of the surface layers
of the ocean during a tropical cyclone. However, in contrast to the results
presented in the preceding paper by Leipper, the observations led to the
tentative conclusion that mechanical stirring was probably more important
than organized upwelling in the cooling process. This conclusion was reached
mainly from the observation that cooling was much more pronounced on the
right hand of the storm track (Figures 1 and 2) where wind and wave action
are known to be most pronounced.

190
REFERENCES

1. Jordan, C. L. and Frank, N. L- 1964 On the Influence of Tropical
Cyclones on the Sea Surface
Temperature Field. Scientific
Report, Department of Meteorology,
Florida State University, 31 pp.

Ba See, a6 We bes 1963 A Study of the Quality of Sea
Water Temperatures Reported in
Logs of Ships' Weather Observation.
Journal of Applied Meteorology, 2,
417-425.

3. -Misher, EB. L. 1958 Hurricanes and the Sea-Surface
Temperature Field. Journal of
Meteorology, 15, 328-333.

191

THE MODIFICATION OF WATER TEMPERATURES
BY HURRICANE CARLA

Robert E. Stevenson and Reed S. Armstrong
Oceanographic Institute
Florida State University

i ie

Rott iain
y UN cial

193
INTRODUCTION

Hurricane Carla entered the Gulf of Mexico through the Yucatan Straits
on September 7, 1961. From there it traveled in a northwesterly direction
and: grew to be one of the five severest hurricanes to invade the Gulf since
1837. By September 10, as it approached the Texas Coast (Figure 1,), pres-
sures in the center were 931.2 mb, and winds of 130 knots whirled around the
eye. Because of the early and continuous advisories issued by the U. S.
Weather Bureau, nearly 500,000 people evacuated the coastal regions. Thus,
despite the fury of the storm and the accompanying storm surges (a maximum
of 7 meters where the storm crossed the coast), few persons were injured.

The energy exchange between the sea and the atmosphere is several orders
of magnitude greater during a hurricane than in less severe tropical cyclones.
Hurricanes provide, therefore, a unique 'laboratory' for investigations of
air-sea interaction. However, the taking of in situ measurements of water
temperature changes is virtually impossible. Aboard ship, nothing can be
done but to practice survival techniques, and even these are unsuccessful on
many occasions. Weather buoys have broken from their moorings, never again
to be seen, and towers have foundered.

It was, therefore, a fortuitous set of circumstances by which changes
in the water temperature distribution affected by hurricane Carla were
"preserved' for investigation at a more peaceful time.

NORTHWEST GULF WATERS IN THE FALL OF 1961

Throughout the year a low-salinity layer of water lies along the coast
of the northwest Gulf. The variations in the salinity, width, and thickness
of the layer are dependent primarily upon the volume of river runoff coming
through the numerous estuaries and lagoons. Usually the surface salinity is
approximately 30.00 per mil close to the shore. A salinity of 36.50 per mil
is normal at distances of 30 to 50 km from the coast. The water of 30.00
per mil may extend to depths of 20 to 30 meters, below which salinities of
36.00 per mil are encountered at 4O to 50 meters.

In September 1961, the brackish surface water lay in a bulge which
extended some 120 km from the coast between the Mexican border and Galveston,
Texas. It was over this bulge that hurricane Carla swept on September 10
and 11 ( Figure 1).

A month later, between October 4-9, scientists cruised aboard the R/V
HIDALGO, of the A&M College of Texas, to investigate the distribution of
temperatures and salinity. Many of the traces from bathythermograph casts
revealed temperature inversions, with magnitudes as great as DSO extending
to depths of 83 meters ( Figure 2). The inversions were all within the area
of the brackish bulge ( Figure 3). Salinities of 29.76 per mil were
encountered near shore, and salinities of 30.00 to 31.00 per mil were measured
though depths of 40 meters. In most of this area, water of 36.00 per mil lay
below 100 meters, although at 70 to 80 meters the salinities were usually

194

Corpus
Christi

SA ue

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11 Sept.

Thermister Tows
23-24 Aug.

The Track of Hurricane Carla on September 10 and 11, 1961,
as Plotted by U. S. Weather Bureau Radar at Corpus Christi
and Galveston, Texas, and Locations of Thermistor-Chain Tows
Made from the R/V HIDALGO

195

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196

Maximum Temperature
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28
| Corpus
Christi

Figure 3. The Distribution of Maximum Temperature Differences in
Inversions as Deduced from Data Gathered on October 4-9,
1961.

197

near 35.00 per mil (Figure A and B). Waters with the steepest salinity
gradient were within 110 km of the shore, and surface waters having salinities
of 33.00 per mil, or less, extended to 210 km from shore (Figure 1A). The
typically steep agian gradients are noted from the curve in Figure B.

The loss of heat in the surface waters to the hurricane atmosphere
lowered the water temperatures, forming the inversions, and, because of the
brackish layer, the lowered temperatures did not result in a density instability
in the water. From climatologic records and previous data collected during
research cruises in the northwest Gulf, it is know that changes in water
temperatures in September and October are negligible. This has been sub-
stantiated by applying the techniques presented by Laevastu (1960) to the
factors influencing heating and cooling of surface waters. Furthermore, a
careful examination of the 91 bathythermograms taken on the October 4-9
cruise showed no indication of heating, or cooling, of the surface waters.

In the weeks following the occurrence of hurricane Carla, winds over the
waters of the northwest Gulf blew at low velocities, and there were several
successive days of light and variable airs. The wind directions were, as
normal, from the southeast. However, the low velocities produced wind-
drift currents of little consequence in re-establishing normal temperature
distribution. Isosteric surfaces, mmputed from data gathered on October 4-9,
were essentially flat. Therefore, no significant density currents were
present.

With these considerations in mind, it was concluded that the water-
temperature structure of the northwest Gulf retained, for at least the suc-
ceeding 4 weeks, the dominant characteristics formed during hurricane Carla.
The inversions which were measured during October 4-9 were believed little
modified from the configuration immediately following the hurricane.

To the southeast of Galveston, where the vertical distribution of salinity
was nearly isohaline, the heat loss from the surface caused instability in
the upper layers. The consequent convective stirring produced an isothermal
layer which extended to depths of 60 meters (Figure 2).

THE TEMPERATURE DISTRIBUTION

The distribution of surface water temperatures in the early days of
October reflected the influence of Carla (Figure 5). Warmer water was
centered in the area where the hurricane deviated from its northwesterly
course, whereas colder water was situated near the outer boundaries of the
low-salinity layer and over that part of the’shallow shelf which was beneath
the track of the storm.

At depths of 50 meters ( Figure 6), the main 'cells' of warm and cold
water were even more sharply defined. The effects of the temperature
inversions were noted where temperatures were slightly greater than 28. 0° oy
which was just more than 0. 5°c warmer than those at the immediately overlying
surface. Farther from shore and to the right of the hurricane track, the

198

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Water Temperatures
Surface
4-9 Oct. 1961

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Figure 5. The Distribution of Surface Water Temperatures on October 4-9,
1961.

200

Water Temperatures
50 meters
4-9 Oct. 1961

28 I
' |Corpus
Christi

‘Figure 6. The Distribution of Water Temperatures at a Depth of 50
Meters on October 4-9, 1961.

201

cooler water indicated modified temperature structures where typical Gulf-
water salinities occurred.

The influence of the hurricane was not restricted to the surface layers,
for there was an upward transport of heat from depths greater than 100 meters.
This was best exemplified by changes which took place in the vertical and
horizontal configuration of the thermocline. For Figures 7 and &, semi-
diagrammatic sketches were drawn from data gathered by thermistor-chain tows
along the tracks indicated in Figure 1. (It must be noted here that
Figures 7 and 8 were drawn from average values of the data obtained by the
thermistor-chain. The temperatures and the depths of the isotherms differ to
some extent, therefore, from those plotted directly from the bathythermograns. )
On August 23 there was a thermocline typical of these Gulf waters, exhibiting
a flat surface at depths of 65 to 70 meters. The thickness of the thermocline
decreased in an offshore diréction, but the 26°C isotherm remained as the
upper limit throughout the length of water measured.

During the hurricane, 26°C water was carried into the disturbed surface
layer, and the 25°C isotherm marked the upper part of the thermocline on
September 15 ( Figure 8). The surface of the thermocline was no longer flat,
having been depressed to the shoreward and seaward of the region where the
tracks of the hurricane and thermistor tow crossed. A comparison of the two
profiles ( Figures 7 and 8) reveals that after the storm the 24° and 25%
isotherms were deeper in the near-shore area and that those below the
thermocline were from 18 to 70 meters shallower than on August 25. The
greater upward displacement was in the deeper water (note the 20°C isotherm,
for example).

The topography of the 25°C isotherm (top of the thermocline) during the
days of October 4-9 closely resembled the configuration of the temperature
distribution. Whereas south of Galveston and seaward of the shelf break,
the thermoclinal surface was generally between 50 and 75 meters, to the west
it was depressed to depths of 102 meters ( Figure 9). From the edge of the
shelf shoreward, the depth to the thermocline decreased abruptly, and, along
the track of the hurricane, it was absent in waters shoaler than 50 meters
(see Figure 2).

Authors' Note: A more complete paper was prepared for presentation at
the Third Intemational Conference on Hurricanes and Tropical Meteorology
held in Mexico City in June 1963. That paper is now in press in Geofisica
Internacional as part of the Conference proceedings.

In presenting this condensed version at Wakulla Springs in February
1965, it was the authors' intent to show the similarities and differences
in the water temperature distribution in the northwest Gulf in 1961 as
compared with that existing after hurricane Hilda (see Leipper, this
volume ) .

202

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204

Depth to Bottom
of Inversions

(mtrs)

Galveston

28
| |Corpus
Christi

Figure 9. The Distribution of Depths to the Base of Temperature Inversions
on October 4-9, 1961.

205
The cruises of the R/V HIDALGO in 1961 were certainly "cruises of

opportunity," for the extent and magnitude of changes in the waters following
a hurricane passage were unknown. Nevertheless, the temperature profile
obtained on September 15, 1961 ( Figure 8), is quite similar to those from
data gathered in 1964. The "stovepipe’ effect beneath the "eye position"
is easily noted, as is the depression of the thermocline on either side.
It would appear that Leipper's analysis is correct; that the warmer surface
water was transported from the areas of the "eye" to lie in "trough" along
the borders of the "eye track," and that the removal of the surface water
beneath the "eye" developed a divergence which resulted in the upwelling
of deeper and cooler water.

One recognizes also, as from the data presented by Leipper, that the
isotherms below the thermocline were at lesser depths following hurricane
Carla than before (compared Figures 7 and 8). This is the case at least to
the depths measured; i.e., about 230 meters. However, the configuration of
the isotherms below 100 meters (perhaps even 75 meters) on September 15,
was nearly identical to that existing before the hurricane. It was strictly
fortuitous that the path of hurricane Carla crossed the southern border of
the western Gulf eddy. Thus, the configuration of the isotherms below the
thermocline existed before the hurricane and clearly was not altered as the
result of any reaction to the storm.

Water temperatures collected in the northwest Gulf in October 1961
differed from those in the north-central Gulf in October 1964. In only the
shallow nearshore waters in the north-central Gulf were there temperature
inversions, whereas deep inversions were widespread in the northwest Gulf
after Carla ( Figure 3). In each case, the inversions occurred only where
low salinity water made up the surface layer.

Hurricane Hilda travelled over Gulf waters of normal salinity (@ 36.00
per mil) throughout most of its course. Hurricane Carla, on the other hand,
travelled over waters which had surface salinities of near 30.00 per mil.
Thus, the water temperatures resulting from the two hurricanes differed sig-
nificantly, and any comparison must be subjective.

It is unlikely that the temperature inversions which were measured in
1961 can have resulted from any mechanism other than cooling; i.e., heat
loss. Upwelling does not produce such vertical temperature distributions.
An introduction of cool, low-salinity water could allow inversions similar
to thase observed to develop. However, the source of low-salinity water is
the estuarine system of the Texas coast which, during the later summer,
contains water with temperatures of 28°-32°c. Furthermore, and as discussed
by Leipper (this volume), the wind field around the hurricanes develops a
wind drift (in these cases, storm surges) which drives water into, rather
than out of, the lagoons and estuaries.

Considering, then, that the inversions represent a certain amount of
heat loss from the waters, the magnitude and depth of the inversions could
be controlled by (1) the thickness and salinity difference of the surface

206

layer, (2) the intensity of the storm and the consequent reaction of the
water, or (3) a combination of both.

Were the temperature decrease sufficient, in any water, to produce a
density instability, convective stirring must take place (in addition to
mechanical stirring by wave action). In such cases, inversions would be
eliminated and a thoroughly mixed layer formed (as noted by Leipper in
waters of normal salinity in the north-central Gulf, and by the authors in
the waters southeast of Galveston). If cooling were insufficient to cause a
convective stirring in the brackish surface water, but extended below the
low-salinity layer, a mixed zone below the inversions would be expected.
None was noted (see Figure 2). A more precise determination could be made
if there were adequate salinity data. However, in October 1961, we obtained
too few salinity samples to analyze the depth distribution of the brackish
water in detail. Thus, we must rely on an interpretation of the temperature
data.

The lesser salinity of the surface waters could, as mentioned, control
the magnitude of the temperature inversions. Again, should the cooling be
so great as to produce a density instability, convective mixing would result.
Conceivably, then, cooling could produce greater temperature differences than
observed in the inversions, but the depth and magnitude of the inversions
would be limited by the consequent instability of the water column.

The temperature-salinity curve in Figure }b shows that, the water to
150 meters (at this station) was far less stable than that normally encountered,
but still did not reach a neutral stability (a frequent condition in the Gulf
during the winter). The temperature decrease was not, therefore, the maximum
possible under the prevailing water conditions. It is clear, then, that the
depth and magnitude of the inversions were the result of the reaction of the
water to the storm.

UPWELLING AND COOLING

The discussion by Leipper of the water temperatures after hurricane
Hilda impressively described the mass transport of surface water from the
region underlying the "eye." Such a picture is not obvious from the data
obtained after hurricane Carla. Rather, the temperature inversions ascribe
to cooling of the water.

Conversely, the temperature distribution after hurricane Hilda
presented great difficulties in defining a degree of heat loss from the
water. One cannot but presume that both mass transport and cooling take
place. However, to fit the two together from the available data is not a
Simple matter, nor have we tried, other than by mental ruminations.

Certainly, were one to place a magnitude of upwelling to the inversions
measured in 1961, then it is clear that cooling extended to a greater depth

207

than indicated by the temperature curves. With the data we have, any
estimation of how mich greater would be pure folly. Nonetheless, it now
seems apparent that the heat loss from the water must have been orders of
magnitude more than the 2.2 x 1018 cai/ok neurs originally calculated.

Acknowledgements - Financial support for this work was from the Office
of Naval Research, under contract NONR 2119(04) NR 083-036.

208

Laevastu, T.

REFERENCES

1960:

Factors affecting the temperature of
the surface layer of the sea, Societas
Scientiarm Fennica, Comm. Physico-Math
XXV 1

209

ON THE LOW LEVEL THERMAL STRATIFICATION OF THE MONSOON AIR OVER
THE ARABIAN SEA AND ITS CONNECTION TO THE WATER TEMPERATURE FIELD

Jose A. Colén
U. S. Weather Bureau
San Juan, Puerto Rico

210

ABSTRACT

During the period August 7 to September 28, 1963 the research vessel
R/V ATLANTIS II, while participating in the program of the International
Indian Ocean Expedition, made several E-W cross sections across the Arabian
Sea carrying out a program of meteorological and oceanographic observations
which included daily radiosondes. This collection of raob data was analyzed
in the form of time and space atmospheric cross sections in order to study
the general properties of the monsoon air.

The southwest monsoon regime is generally established over the Arabian
Sea around the latter part of May and is maintained as a fairly steady and
persistent current until the latter part of September. Therefore, the data
collected by the ATLANTIS II showed characteristic properties of the monsoon
current at the height of the season.

The thermal structure consisted essentially of two layers of air: a
shallow layer of humid air near the surface and a dry, relatively unstable
air mass on top, separated by a pronounced thermal inversion. The nature
of the monsoon inversion and of its distribution over the Arabian Sea was
investigated.

The properties of the surface moist layer and the modifications during
its path downstream will be discussed also in relation to the water tempera-
ture field and the weather-producing processes responsible for monsoon rains.

ark
INTRODUCTION

This report deals with the low level thermal stratification of the

' atmosphere over the Arabian Sea during the summer monsoon season, its
relation to the water temperature field, and to the general atmospheric
circulation in the area. In order to understand better the problems involved,
we should review briefly the general characteristics of the Indian monsoon
circulation, although these are fairly well know to most readers.

The mean circulation near the surface over the Arabian Sea, once the
southwest monsoon current is well established, can be illustrated by the
mean chart for August (Figure 1). The monsoon circulation is usually
established over the northern Indian Ocean by early June. The flow pattern
shown in Figure 1 persists with little variation from June until late
September. The surface flow over the Indian Ocean northward from about
latitude 20°S consists essentially of a current formed by the southeast trades
of the southern Indian Ocean turning in a clockwise direction near the equator
to continue as a broad southwest current over the Arabian Sea into the Indian
subcontinent and southeast Asia.

During the 1963 season we had the opportunity to analyze and study the
Indian monsoon weather while stationed in Bombay participating in the
activities of the International Indian Ocean Expedition. One of the features
noted with great interest was the persistent and steady character of the
circulation over the Arabian Sea which showed very little interdiurnal varia-
tions in a picture that day after day differed little from what appears in
Figure 1. The daily charts showed important variations in wind speed, but
only small variations in the direction of flow.

The establishment of the southwest monsoon current over the Arabian Sea
brings about significant change in the oceanic circulation and distribution
of water temperatures. There is as a result a large variety of air-sea y
interactions going on over this oceanic area during the monsoon, with signifi- Vv
eant effects on both the properties of the oceanic surface layers and of the
atmospheric layer above. Some of these effects were studied in an earlier
report (Colén, 1964). The Arabian Sea during the summer season presents as
varied and striking evidence of air-sea interactions as can be found anywhere
else in the globe. Some of these developments assume even greater importance
when considered in the light of the vast production of rain which normally
takes place a little farther downstream over India.

: Among the activities carried out during 1963 in Bombay were a series
of aircraft flights over the Arabian Sea by the U. S. Weather Bureau-Research
Flight Facility and by the Woods Hole aircraft operated by Dr. Andrew Bunker
to investigate the monsoon flow over the ocean. One important reconnaissance
mission consisted of a flight upstream following the surface streamlines,
which in a period of 2 days covered the path all the way from Bombay to the

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equator near the coast of Africa. That mission contributed greatly to our
knowledge and understanding of the atmospheric properties and weather
features over that oceanic body. Another important source of data were

the cruises by the research vessels RV ATLANTIS IT and RV ANTON BRUUN. The
ATLANTIS II carried out an extensive survey over most of the Arabian Sea
during the period August 7 - September 28, 1963.

The present study is mainly concerned with some aspects of the raob data
collected by the ATLANTIS II. A collection of about 45 soundings made aboard
the ATLANTIS II were studied and analyzed in various forms which revealed
well the general characteristics of the thermal stratification over the ocean.
We also had access to 10-12 soundings made aboard the RV ANTON BRUUN from
August 11-26, 1963, and to innumerable dropsondes made by the research
aircraft.

TRACK OF THE ATLANTIS II IN RELATION TO THE FLOW AND
WATER TEMPERATURE FIELDS

; The track of the ATLANTIS II in relation to the monsoon circulation is
illustrated in Figure 1. The monsoon current shows a significant velocity
maximum in the western side of the sea near the coast of Somali, with speeds
decreasing downstream. In the mean picture the velocity maximum is around
30-35 knots, but velocities of around 50 knots were measured by aircraft in
that area. Very significant upwelling of the cold subsurface water is also
observed in that area.

One interesting and significant feature of the flow is that nearly all
the surface air that enters into India seems to have an oceanic source
originating in the Southern Hemisphere. However, flow charts at some distance
above the surface, for example 850 mb, reveal flow off the African Continent
and the Arabian peninsula moving eastward toward India. Thus, the southwest v
oceanic current is extremely shallow.

The ATLANTIS II made four latitudinal cross sections between the east

and west sides of the sea. She cruised from Aden to Bombay along latitude
15°N from August 6 to 15. From Bombay she cruised westward along latitude

20°N and southward to the region of upwelling near Somali; then eastward along
latitude 10°N, arriving in Colombo, Ceylon on September 7. From Colombo she
proceeded westward along latitude SON, to the coast of Africa, then went south-
ward arriving in Zanzibar on September 28.

This track turned out to be quite good from the point of view of a study
of the monsoon current. The latitudinal paths were to large extent across
the flow in the western Arabian Sea, but nearly parallel to it in the east
side. The track from August 23 to 30 followed for a distance of about 1000
miles a direction upstream closely parallel to the flow, terminating in the
cold-water region near the coast of Somali.

214

The distribution of water temperatures is shown in Figure 2. The
isotherms are oriented in a general southwest-northeast direction with very
cold temperatures of 22-23 C in the western coastal areas in the regions of
upwelling and warm centers of 27-28°C in the east side. The influence of the
atmospheric circulation on the distribution of temperatures can be easily
visualized. We can note with interest the extremely warm temperatures in
the Gulf of Aden and the thermal gradient between the Gulf and the Arabian
Sea. The track of the ATLANTIS II is also reproduced in Figure 2 to show the
distribution of data with respect to the water temperature field.

THERMAL STRATIFICATION - THE MONSOON INVERSION

Two soundings obtained by the ATLANTIS II in the central Arabian Sea,
reproduced in Figure 3, illustrate the essential characteristic of the thermal
stratification of the monsoon air over the ocean. One was obtained on
azo Ii, 1963, near 15 N, and 58°R, and the other on August 14, 1963, near
16-N, and 63°E. Only the dew-point curve for August 14 is illustrated.

Figure 3 indicates a well-mixed, humid layer of air near the surface, a
pronounced inversion immediately above in the levels from about 900 to 800 mb: ,
and a relatively warm, dry-air mass aloft. The lapse rate is close to dry
adiabatic in the surface layer, stable in the inversion, relatively unstable
from the top of the inversion to the 500-mb level and close to the moist adia-
batic in the high troposphere. The humidity is large in the surface layer and
drops off significantly at the base of the imversion. These data illustrate a
moist air mass below, evidently dominated by oceanic influences and a dry hot-
air mass aloft, presumably of continental origin. Figure 3 shows a larger
depth of the moist layer, or higher base of the inversion, in the position
farther east.

The presence of such a pronounced inversion so close upstream from the
coast of India was detected quite early in our studies of the monsoon air over
the ocean and presented some interesting questions concerning the mechanisms
for development of monsoon rains.

The soundings were analyzed in the form of five cross sections; four of
them along latitudinal directions and one along a direction parallel to the

' flow. They all presented a consistent picture of a low inversion in the

western edge of the sea, which rose eastward toward India and showed a
pronounced tendency for dissolution near the Indian coast. The cross section
along latitude 15N showed a rather deep stable layer in the western end with
a warm center of over 20°C temperatures at the 825 mb level, which was about
8°C warmer than at the same level in the east side of the ocean. The top

of the moist layer or base of the inversion rose from the 960 mb level at
longitude 50 & to the 850 mb level at longitude 68 E. A sounding obtained
near the Indian coast on August 15 showed ng presence of the inversion. The
other E-W cross sections along latitudes 10 N and 5 showed more or less
similar characteristics.

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The cross section made from August 23-30, which extended in a line
closely parallel to the direction of flow and indicates well the modifications
introduced as the air moved downstream over the sea_surface ig illustrated
in Figures 4 to 6. It starts in a position near 10 N, and 57 E, very close
to the,center of colder temperatures, and extends to a position near 20 N,
and 63 E. Figure 4 shows the temperature field. The base of the inversion

or top of the moist layer appears at the 960 mb level in the upstream end;
it rises slightly in the first 600 miles, and more rapidly afterward. In
the downstream end it was observed at the 850 mb level. The depth of the
stable layer did not vary much gownstrgam. At the surface there was an
increase in temperature from 24 to 26 C between the upstream and dowmstream
end, which no doubt resulted from heating from the water surface. At the

850 mb leyel the temperature decreased from 20-21 C in the upstream end to
around 16°C in the northeastern or downstream end.

The field of potential temperature, Figure 5, indicates as one of its
most interesting features that the 300°K line followed closely the level of
the base of the inversion. The 304 K isoline also followed closely varia-
tions in height of the inversion layer. This is evidence of rising adiabatic
motion as the monsoon air moved downstream over the ocean, since in the lower
layer we can presume with a great degree of validity that the flow followed
closely along the cross section line. At upper levels it cannot be similarly
assumed that the flow followed parallel to the cross section. The potential
temperature lines above the inversion were displaced generally downward
downstream. An attempt to construct an isentropic chart for a level in the
dry air mass aloft was not too successful, but there appeared to be a tendency
for downslope motion between the west and east sides of the sea, a tendency
that is also supported by the data in Figure 5.

The most important feature of the moisture distribution ( Figure 6)
is the increase in moisture content downstream near the surface, evidence
also of the exchange processes between the sea surface and the air above.
The tendency for higher moisture content in the upper levels near 500 mb in
the eastern end of the section is also evident. In computations carried _gut
over_the central Arabian Sea average evaporation rates of 600-700 cal em
day~2 were obtained.

An analysis of the distribution of the height of the base of the inversion
or mean depth of the surface moist layer is shown in Figure 7. It indicates
a shallow moist layer or low inversion in the western section of the Arabian
Sea, specially near the coasts of.Arabia and Somali, and a rise eastward
toward the coast of India.

In the western half of the Arabian Sea all available soundings revealed
the presence of the inversion near the surface. A few aircraft soundings
available in the northern corner near Arabia and Pakistan revealed a very
pronounced inversion close to the surface. In the eastern third, near the
coast of India, most of the available observations showed absence of the
inversion. The stratification was generally indicative of disturbed weather
conditions with lapse rates close to the moist adiabatic and high moisture
content extending to 500 mb and above. To the south, near the equator, the

218

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Figure 4. Temperature (°C) cross section based on data collected by RV
ATLANTIS II between positions 20°N, 63.5°E, and 10°N, 52.5 E,
August 23-30, 1963. Heavy solid line shows base and heavy
dashed line top thermal inversion.

219

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500 — oan —
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Figure 5. Potential temperature (°K) cross section for section described
in Figure 4.

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eross section taken along latitude SN, showed the presence of the inversion
in most of the soundings, generally higher in the eastern section. In this
eross section and also in the one taken along latitude 10%, the presence of
a second inversion near 600 mb was noticed in most of the soundings.

In the western equatorial region near the coast of Africa conditions
were more variable. Generally, more unstable conditions prevailed there.
Most of the soundings in that area showed moisture extending to high levels.
A few soundings obtained from the research aircraft earlier in the season
showed tendency for an inversion near 850 mb.

In general, the distribution showed a pronounced low inversion over the
cool waters in the upwelling areas near the coasts of Somali and Arabia, a
general rise in the inversion downstream toward the coast of India and a
pronounced tendency for sudden and violent overturning with disturbed weather
in the coastal area just west of the Indian coast. The disturbed weather
zone along the east coast extended south from about latitude: 20°; north of
this latitude more stable and drier conditions were prevalent, which is of
course reflected in the presence of the desert-like conditions of northwest
India and west Pakistan.

SUMMARY AND CONCLUSIONS

Some remarks should be made concerning the basic assumption behind the
method of analysis used. The data were all obtained from a moving platform
and in the case of ATLANTIS data covered a period of 7 weeks. The aircraft
data mentioned was obtained about a month and a half earlier. They were
analyzed to show variations in space, under the assumption that the variations
in time were negligible compared to those in space. On the basis of the
experience during the 1963 season, we think that this principle holds well.
Of course, not all of the details shown by the data can be considered
indicative of space variations exclusively. The gross features emphasized
above are unquestionably evidence of space variations. The validity of this
approach is supported by the fact that the picture derived from the analysis
is quite close to the one to be expected considering the factors operating
over that area.

The picture can be summarized as follows: The establishment of the
monsoon circulation near the surface over the Arabian Sea in the characteristic
manner illustrated in Figure 1 and in reference to the land masses of Africa
and Arabia influences greatly the distribution of water temperatures over the
Arabian Sea mainly through the processes of upwelling in the western side
of the sea. At the same time the oceanic influence on the air mass in the
surface layer combined with a continental air mass that moves eastward in
the levels upward from about the 900 mb level results in stabilizing the
thermal stratification in the western zone of the sea. We have thus a
relatively cool moist shallow air mass below and a dry warm mass aloft
separated by a thermal inversion and moisture discontinuity. These

stability conditions in the west central Arabian Sea seem to persist with a
minimum of variation throughout the monsoon season.

As the monsoon current moves northeastward over the sea, flux from the
sea surface, mixing enhanced by the relatively strong flow near the surface,
and adiabatic ascent imposed by the field of motion act to increase the heat
and moisture content of the monsoon air and to increase upward the depth of
the moist layer. The flux of heat from the sea_to the ocean in this area has
been computed as close to 600-700 cal cm™ cava At the same time sinking
motion in the warm air mass aloft contributes also to vertical mixing with
the moist mass below.

In the western two-thirds of the sea the properties of the monsoon J
current appear to be influenced almost exclusively by the process of air-sea
property exchange. In the eastern third of the oceanic mass, closer to
the Indian coast, the prevailing tendency is for dissolution of the inversion,
violent vertical mixing leading to generally unstable conditions with
considerable layered cloudiness and convection. The prevalence of unsettled
weather in that area throughout the monsoon season is very well supported by
the mean weather distribution charts. The destruction of the inversion cannot
be ascribed to purely air-sea interaction processes; it is apparently due to
developments in the synoptic scale. Conditions in the area to the west of
Bombay suffer significant interdiurnal variations and periods of relatively
good weather alternate with periods of more violent and widespread rain
activity. Periods of so-called breaks in the monsoon over the west coastal
sector of India seem to be associated with conditions over the Arabian Sea
such that the relatively stable conditions over the ocean are extended east-
ward into the Indian coast. Our studies in Bombay indicated synoptic develop-
ments at 500 mb as being largely responsible for weather variations along
the west coast of India; the low level flow showed little evidence of the
synoptic systems.

The conditions of flow and the distribution of the water and land masses
are such that the air that penetrates the west coast of India properly has
had considerable oceanic influence and has great rain producing potential.

On the other hand on account of the decrease of the water mass northward and
associated increase in the influence of the Arabian land mass, the air mass
moving into west Pakistan and northwest India, even at the height of the
monsoon season, has a minimum of oceanic influence. This is an important
contributing factor to the heat and dryness that prevails in that area.

As mentioned at the beginning ,the Arabian Sea during the southwest
monsoon season offers many interesting examples of air-sea interactions, and
every corner of the ocean exhibits conditions that differ significantly from
the others. There is no question that the processes of air-sea interaction
play a major role on rain-producing processes over the Indian Subcontinent.

22h

REFERENCES

ne Jose A. 1964 On Interactions between the Southwest
Monsoon Current and the Sea Surface
Over the Arabian Sea; Indian Journal

of Meteorology and Geophysics, 15,
183-200.

225

A LOW LEVEL JET PRODUCED BY AIR, SEA AND LAND INTERACTIONS

Andrew F. Bunker
Woods Hole Oceanographic Institution
Woods Hole, Massachusetts

226

ABSTRACT

The summer monsoon blowing from the southwest over the Arabian Sea
produces upwelling of very cold water off the coasts of Somalia and Arabia.
Aircraft meteorological observations show that a narrow, 25 msec jet

/ exists at about 600 m over the cold water off Somalia. Analysis of data
available at the moment indicates that this jet results from the combina-
tion of reduced frictional drag of warm air over cold water and the thermal
wind shears produced by the cooling of the lower air by the water.

227

ABSTRACT

As the southwest monsoon develops over the Arabian Sea, a sequence of
sea, air, and land interactions occurs that culminates in the production of V7
a 25 m sec. low level jet at 600 to 1000 meters blowing off the coast of
Somalia. Aircraft and surface observations are presented which describe the
jet, the thermal structure of the atmosphere, the sea surface temperatures.
and the surface pressures. From these data and climatological data it is
shown that the jet results from a sequence of two complete cycles of thermal
reactions of the air to sea and land temperatures and kinetic reactions of
the water to atmospheric wind stress. These cycles occur on decreasing size
scales but with increasing intensities thereby producing an intense local jet.
The geographical location of the jet is fixed by the configuration of the land
masses and proximity to the equator. The existence, general shape. magnitude.
and geographical position of the jet are explained as resulting from (a) a
large land mass north of the Indian Ocean, (b) = land mass to the west of the
Ocean, (c) strong heating of the land which intensifies the pressure gradient.
(ad) a small value of the Coriolis force, and (e) air-sea interactions which
produce through upwelling of cold water and cooling of the air thermal winds wi
and a low frictional drag of the air over the water.

INTRODUCTION

As part of the International Indian Ocean Expedition, meteorological
studies were carried out over the Arabian Sea using a C-54Q aircraft. This
airplane was equipped with instruments designed to measure temperatures,
winds, humidities, clouds, radiation, and turbulence. On August 30, 1964. a
flight was made from Aden cutting perpendicularly across the wind blowing from
‘the SW off the coast of Somalia. The track outbound was flown at 100 to 600
meters. A climb was made to 4500 m at 4° N, 56° E. The return flight to
Aden was made at 4500 m and 5 radiosondes were released enroute. On
September 1, 1964, a track was flown from Aden to Bombay cutting the Somali
jet again. The portion of the track from Aden to 11° N, 58° E was flow at
low levels while the remainder was flown at 4500 meters. From these observa-
tions it becomes clear that the strong_yinds take the form of a low-level jet
with a maximum speed of about 25 m sec

DESCRIPTION OF THE JET

Figures 1 and 2 are presented to show the main characteristics of the
wind system as it existed on the 2 days, August 30 and September 1, 1964.
The wind speed - height curve, Figure l, was drafted from the Doppler radar
winds observed during the aircraft's ascent at about 11 N, 58 E. It is
seen that a large wind shear exists in the 500 to 1000 meter region, that a
rather broad maximum exists in the 1000 to 1500 meter range, and a steep
negative gradient exists above this level.

228

DOPPLER RADAR WINDS
4000

220° e 1 SEPTEMBER 1964

11N 58 E
3000 :

2500

2000 e

HEIGHT, METERS

1500 :

1000

500 .

0 5 10 15 20 25 30
WIND SPEED, M/SEC

Figure 1. Doppler radar winds obtained from the C-54Q aircraft are
plotted against height.

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Figure 2 presents the horizontal distribution of the observed wind
vectors on the same 2 days. On August 30, the aircraft crossed the axis of
the jet at 10° N, 530 E at 550 meters. At this point the velocity was about
25 msec. One hundred km either side of the axis the velocity drops to
about 15 m sec” indicating a rather sharp narrow jet.

On September 1, 1964, a track about 45° to the wind was.flowm. The plot
of this flight shows that a strong narrow jet with 27 m sec winds was
encountered in the vicinity of le N, Bae E. Beyond this area the wind_dropped
off to speeds of about 16 m sec-l. As the ascent was made at ae nq, Sis) 18;
the aircraft passed through a jet at 1000 meters, with 25 msec speeds. One
would like to know whether this is the same jet that was encountered at 12° N
53° E ata neu iiate of 560 m. Tt is possible that it is a broad jet extending
mom IZ? iy, SO im wo We iy 58° E with a level of the maximum wind that rises
in the southeasterly direction. High values of the turbulence encountered
along this track suggest that the aircraft probably was flying underneath a
wind maximum. This point will be discussed later after the presentation of
the turbulence data.

It should be stated at this point that the jet is not a transient
phenomenon that happened to exist in the area on the days that the flights
were made. Rather, the intense winds in this area are a persistent feature
of the southwest monsoon, as a casual inspection of weather charts of the area
will reveal. Ship reports show relatively minor variations in strength and
position from day to day.

THE SEQUENCE OF INTERACTIONS BETWEEN AIR, SEA, AND LAND
AND THE DEVELOPMENT OF THE SOMALI JET

WA (1) Global-Seale interactions. The first phase in the development of
the wind system occurs in the northern hemisphere in the spring on a scale
involving the entire continent of Asia and the Indian Ocean. With the returm
of the sun to the hemisphere the land masses of Asia and North Africa are
warmed. The land in turn warms the atmosphere and gradually changes the cold
highs to warm lows. During this same time interval the temperature of the
Indian Ocean south of the equator remains about the same or cools a small
amount. The maintenance of the water temperature keeps the air temperature
about the same and hence the surface pressure remains the same or increases a
bit. By May the pressure in the north has been reduced to about 1005 millibars
while the pressures in the south have increased to about 1022 mb: This pres-
sure gradient accelerates the sir mass northward across the equator and north-
easterly across the Arabian Sea.

Leas Oceanic-Scale interactions. The reaction of the sea water to this
moderate air flow from the south is the movement of the surface water to the
east away from the coast of Africa. The presence of the African Continent
prevents replacement of the water except by upwelling of deeper cool water.

As a result of this replacement, the waters from the Mozambique Channel to
north of the equator become a few degrees cooler than the water 2000 kilometers

231

to the eastward. The effect of this cool water is to cool the lower atmos- /

phere and thereby increase the surface pressure. Normally at this season a
belt of high pressure lies along the 10° 5 latitude line at 500 millibars.
With the cooling of the lower air below and to the north of this belt, the

surface high develops a weak ridge extending northward along the African
Coastal waters to about 10 N. The characteristics and semi-permanent nature
of this ridge of high pressure are clearly shown on the IGY Tropical Zone
Weather Maps published by the Seewetteramt, Deutscher Wetterdienst. Hamburg.
Study of the maps for July, August, and September 1957, shows the ridge to
be present on all maps with minor variations in amplitude and position.
Adjacent to this ridge is a trough of low pressure lying to the west over
Kast Africa. This trough is caused by the intense solar heating of the land
and air. Its effect is_to greatly increase the pressure gradient in the

region around 10 N, 50 E. The surface pressure map for August 30, 1964,

has been drawn from data made available by the International Meteorological
Center, Bombay, and presented as Figure 3.

“(3) Local interactions. As a result of the steep pressure gradient
developed along the coast of northeast "frica, the air that has moved slowly
across the equator into the regions begins to accelerate and quickly attains
higher velocities. Once the air has attained relatively high velocities
north of about 5 N, the transport imposed upon the surface water is greatly
increased. With this increase in transport the upwelling of bottom water is
greatly increased and the surface temperatures drop many degrees over an area
of more than 100 square degrees. Figure 4 presents a chart of surface water
temperatures drawn by H. Stommel and B. Warren of W.H.O.IT. The data were
obtained from the research vessels, ARGO, of Scripps Institute of Oceanography
and DISCOVERY. of the National Institute of Oceanography during their cruises
in August 1964, tothe Somali Current region. It is seen that immediately,
off the coast at 9 N, exceedingly cold water with temperatures down to 13 @,

was observed.

This colder water continues the interaction cycle by cooling the sir at
a greatly accelerated rate. The intensity of this cooling is shown in
Figure 5 which presents dropsonde and psychrograph data obtained from the
c-54 aircraft on August 30, 1964. The figure is a cross section of the
atmosphere from the Gulf of Aden to 4° nN, 56° E. Potential temperatures
were plotted on the diagram and isentropes drawn from the data points. The
cooling of the air by the water about 200 kilometers southeast of the mouth
of the Gulf of Aden is very apparent. This highly localized cooling is
superimposed upon the pattern of a general warming of the air from south to
north.

To explain the increase in wind speed with height to the jet maximum
and its subsequent decrease above this level, the temperature cross section
made on August 30, 1964, must be studied. It will be noted that the tempera-
ture increases to the right of the wind, indicating that the wind will in-
crease with height. Above the 900 mb level or the level of the jet maximum,
the air temperature decreases to the right of the wind and hence the wind
should decrease with height. Measurement of the horizontal temperature
gradient and application of the thermal wind shows that the wind should

232

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008

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006

008

174

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bp96t
LSNSNV OF

40° 45° 50°

233

SEYCHEL

LES oF

40° : 45° 50°

5°

Figure 4. Surface water temperature chart constructed from ARGO and

DISCOVERY data by Stommel and Warren.

23h

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235

increase about 10 m sec : through the lowest 600 meters. As this situation
is one in which the Coriolis and pressure gradient terms of the equation

of motion are not balanced, the increase in wind with height probably does not
attain the full geostrophic value. The wind speed increase from 100 m to
1000 m on September 1, 1964. was about 12 m sec. which is in fair agreement
with the computed thermal wind increase.

One additional bit of information is now presented that will aid in the
analysis of the generation of the low level jet. The magnitudes of the root-
mean-square turbulent vertical velocities of the air have been measured from

the aircraft's accelerometer according to the method of Bunker (1955) and
expressed_in cm gec7!. The observations were made on September 1, 196k,

around 11° N, 57 E. Also plotted on the height-vertical velocity diagram,
Figure 6, are values of the turbulent velocities observed in other regions

of the world. Several features of these data are very unusual end significant.
First, it is noted that the turbulent velocities at the lowest level are about
the same as the turbulent velocities measured in the trade winds of the Worth
Atlantic Ocean by Bunker (1955). From this observation of weak turbulence it

is concluded that the frictional drag of the water, on the air is nearly the

same as the drag found under the lighter (5 m sec ) trade wind situation and we
is therefore _many times smaller than would be expected for high winds of 15

to 20 m sec”

The small frictional force coupled with the small Coriolis force, due to
the small Coriolis paramater near the equator, cannot balance the pressure
gradient force and hence the air particles are accelerated rapidly across the
isobars. In the region of the strongest winds, it appears that geostrophic
balance is never attained. Down wind of the jet the pressure gradient
decreases and the wind becomes geostrophic. To understand and prove these
relationships in this region a much more detailed and quantitative study of
the terms of the equation of motion must be made.

The second noteworthy feature of Figure 6, is the high turbulent velocities

observed at the 600 meter level. For comparison, observations made at various
heights in a strong wind (20 m sec.) Situation over the North Atlantic Ocean

by Bunker (1960) are plotted on the diagram. It is seen that after a rapid
increase in turbulence in the lowest layers the turbulence decreases rapidly
with height. Such a turbulence-height curve is characteristic of a situation

in which the turbulence is generated by the flow of air over a surface and
decays aloft. In the present case it appears that only a small amount of
turbulence is generated by the flow of the stable air over the water and that V
a greater amount is generated at a higher level where the air is not as stable.
It is concluded from this trace that the turbulence must be generated by high
wind shears and that the aircraft was flying below the level of the maximum

winds. If this is true then ghe jet observed at 12 N, 53° E, at 500 m
probably extends to 11° N, 58° E, where it was observed at 1000 meters.

236

HEIGHT, METERS

TURBULENCE — HEIGHT DIAGRAM -

SOMALI JET @ 1 SEPT.1964
ATLANTIC TRADE WINDS X MARCH & APRIL 1953
N. ATLANTIC WESTERLIES © 14 JAN.1955 (20M/SEC)

700

600

500

400

300

200

100

O 20 40 60 20 100 120 140

TURBULENT VERTICAL VELOCITY, @,, CM/SEC
Figure 6. Turbulent vertical velocity values plotted against height.

237

SUMMARY

The observations made off the coast of Somalia describe the low level
jet system. One question is left open concerning its geographical limits.
It is possible that the system consists of a single jet broadening downwind
and varying in height across stream. Another possibility is that the system
consists of more than one jet.

Interpretation of the observations leads to the conclusion that the jet
is formed by a series of interactions between the land. sea. and air. Also ‘7.
it is evident that the jet depends upon the configuration of the land masses
and their proximity to the equator. It is this dependence on position and
configuration that restricts the jet to a relatively small area off the coast
of Somalia and makes it a semi-permanent feature of the southwest monsoon.

Acknowledgement -

The research described in this paper was supported by National Science
Foundation Grant 22389. The C-54Q aircraft was bailed to the Woods Hole
Oceanographic Institution by the U. S. Navy through the Office of Naval
Research.

238

Bunker, A.F.

Bunker. A.F.

REFERENCES

UND) =

1960 -

Turbulence and shearing stresses
measured over the North Atlantic
Ocean -by an airplane-acceleration

technique. Journ. Meteor., Vol. 12,
WY5-455.

Heat and water vapor fluxes in air
flowing southward over the western
North Atlantic Ocean. Journ.

Meteor., Vol. 17, 52-63.

U. S. FLEET NUMERICAL WEATHER FACILITY ACTIVITIES RELATING
TO SEA-ATR INTERACTIONS ON A SYNOPTIC SCALE

Cdr. W. E. Hubert, USN

Fleet Numerical Weather Facility
Monterey, California

"The opinions presented in this paper are those of the author
and do not necessarily represent the official views of the
Navy Department at large.”

239

2h0

ABSTRACT

The mission of the Fleet Numerical Weather Facility is to provide
numerical weather products on an operational basis peculiar to the needs
of the Naval Establishment and to develop and test numerical techniques in
meteorology and oceanography applicable to Naval Weather Service analysis
and forecasting problems. At Monterey the method of approach to these as-
signments has been to treat the atmosphere and the oceans as one environ-
ment with particular attention being directed toward interactions between
the two media which constitute this environment.

The purpose of this paper is to summarize the various analysis and fore-
cast programs currently utilized at Fleet Numerical Weather Focility
Monterey and to outline future plans. Those programs which depend more heavily
on quantitative computations of air-sea interaction or on strictly maritime
meteorological observations are discussed in more detail. These include:
sea surface temperature analyses (and their scale and pattern separation),
sea and swell analyses and forecasts, and synoptic current analyses (and
their use in estimation of convergence in the sea and their effects on
thermal structure).

Finally, a description will be given of the master scheme which is being
developed for numerical analysis and prediction of oceanographic elements.

2h1
INTRODUCTION

One of the assigned missions of the U. S. Fleet Numerical Weather
Facility (FNWF) at Monterey, California, is to prepare meteorological and
oceanographic analyses and forecasts in support of fleet and other operations
throughout the Navy. As the activity's title implies, these products are
prepared numerically using the latest high-speed electronic computers. The
approach used at FNWF has been to apply a combination of dynamic theory and
empirical experience to problem solving by computer. In general, only

problems which have direct Navy application and which show promise of opera-
tional usefulness within l-year's time are undertaken at FNWF. In this
sense, the developmental efforts at the facility should be called "applied"
rather than "basic research."

While early efforts at Monterey were concentrated ‘on atmospheric
analysis and forecasting, emphasis has been shifting more and more in the
last 2 years to oceanographic problems and, in particular, to sea-air
interactions. The atmosphere and the oceans are considered to be one
environment as far as naval operations are concerned. Each of the media
affects conditions in the other and their behavior should not, and cannot,
be treated independently. The development of improved environmental analyses
and forecasts on a synoptic basis demands that we account for, in a quantita-
tive manner, the energy exchanged between sea and the atmosphere.

NAVY PROBLEM AREAS INVOLVING SEA-ATR INTERACTION

While the entire problem of sea-air interaction involves the transport
of some property (momentum, heat, moisture, etc.) between the two media, the
principal effects now being studied at Monterey can be broken down into sub-
areas. One general class of programs involves the transfer of atmospheric
momentum to the sea (generation of sea,swell and surface currents), and the
other primarily deals with heat exchange at the air-sea interface.

Heat exchange obviously works in both directions. Transfer to and
from the atmosphere must be included in atmospheric forecast models. The
relationships between synoptic scale heat exchange and weather now under
investigation at FNWF will be covered by Dr. Laevastu (1965) later in this
conference. The exchanges which influence sea surface temperatures (and
consequently subsurface thermal structure) are of primary interest in
Anti-Submarine Warfare (ASW) applications and will be covered in further
detail herein.

Wind-driven waves (and swell) are important from several viewpoints.
They not only affect day-to-day operations but can be critical in underway
replenishment, launch and recovery from a carrier, optimum ship routing,
amphibious operations, etc. In ASW problems they influence sea surface
temperature, layer depth and thermocline intensity through mechanical mixing.

ake :

Wind-driven and thermal currents are of minor importance to naviation,
search and rescue but play a major role in modification of layer depth through
large-scale mass convergence/divergence in the surface layers of the oceans.
In the eastern North Pacific, convergence/divergence considerations have
frequently been found to outweigh all other terms contributing to changes in
layer depth.

Heat transfer is of primary interest to ASW operations in that surface
cooling causes convective mixing (resulting in a deepening of the mixed
layer and intensification of thermocline gradient) while surface heating
forms a shallow, transient surface thermocline (in the absence of mechanical
mixing through wave action).

All of the above processes of interaction either contribute to sea
"noise" or modify surface and subsurface temperature structure, and therefore
influence sound propagation in the oceans. An accurate analysis and/or fore-
east of the state of the total environment as influenced by sea-air interaction
is the key to successful naval operations in general and to the Anti-Submarine
Werfare Environmental Prediction System (ASWEPS) in particular.

The following table summarizes the principal environmental inputs to
ASWEPS; there are others, such as bottom effects, but those listed here are
the problem areas under attack at Monterey(see Table I).

THE FNWF MASTER SCHEME FOR OCEANOGRAPHIC ANALYSIS AND PREDICTION

Figure 1 outlines the master scheme developed at Monterey for numerical
analysis and prediction of oceanographic elements and processes (see FNWF
Tech. Memo No. 5). If nothing else, the figure shows the complexity of the
numerical program being undertaken by FNWF in the general field of interest
to this conference.

Looking first at the colum headed Basic Data, one can see that a large
part of the input data to this program is derived from meteorological observa-
tions. Since the number of BT observations is insufficient for truly synoptic
oceanographic analyses (except perhaps in limited areas), we are forced to
derive the maximum information from meteorological reports at the ocean surface.
The basic approach at FNWF has been to obtainthe first estimate of oceanic
thermal structure from purely exchange considerations and then to modify this
"suess'' with BT data where available.

The various computations involved in this method of oceanographic analysis
and prediction are summarized in the middle of Figure 1 under the column headed ©
Computed Quantitites and Processes. The flow diagramming leading to and result-
ing from these computations serves to emphasize the entire concept of sea/air
exchange utilized in the FNWF oceanographic scheme. Details of some of the
more important components of this overall plan will be covered herein or by
Dr. Laevastu.

2h3

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/SONHOUAANOD SNTd DNIXIN HAILOMANOO CNV IVOINVHOSW WOM GHAIYHC -

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TVOLOOTOMOU“LEN ATIYVNIYd ONTATOANT HVINNYOL GONVHOXA LVHH WOU CHLAdMWOD -

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FLEET NUMERICAL WEATHER FACILITY

SCHEME OF NUMERICAL ANALYSES AND PREDICTION
OCEANOGRAPHIC ELEMENTS AND PROCESSES

FLOW, COMPUTED REPROCESSING | AND NAL
WNEUT FROM) Baste | OBR PROCESSING (QUANTITIES INTERDICIPLINABY PRODLICTS _
FORECASTS | CaseRveD. ano cuimato- | AND USE AND PROCESSES US AND OUTFUTS
SYNOPTIc. LOGICAL DATA.

a ba
TEMPERATURE SNAWSES ry

SEA ie ade
SURFACE lee
TEMPERATURE

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RELATIVE
HUMIDITY

CLOUD
COVER, IN
3 LAYERS

bay? Soran

DAY,
ALTITUDE

TOTAL HEAT EXCHANGE
FORECAST, 24, 46, UP.

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FORECAST, 30,48, HA.

CONVECTIVE}
STARING

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TWeRMaL, STRUCTUBE [41
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FEEDBACK TO
MET. CASTS
5 Av

Figure 1. Fleet Numerical Weather Facility, Monterey Master Scheme
of Numerical Analysis and Prediction of Oceanographic
Elements and Processes.

2h5
EXAMPLE OF FNWF OCEANOGRAPHIC PRODUCTS

A. SEA SURFACE TEMPERATURE

The FNWF Sea Surface Temperature (SST) analysis program uses as input
data all ship injection temperature reports from the previous 3 1/2 days. A
median seeking technique is utilized to reduce the influence of erroneous
reports. All observations within C.7 mesh lengths (about 150 miles) of each
erid point are compared with the previously analyzed value at each ship's
position. If a new observation is warmer than the previous analysis, the
value at the nearest grid point is raised a fixed amount (0.1C); similarly,
the value is lowered the same amount if the observation is colder. Areas of
no data are modified only by relaxation and smoothing.

Figure 2 is an example of a numerical analysis of this parameter.
Details of this program and error distributions of SST observations have been
described by Wolff (1964). The SST analyses made at FNWF, Monterey are
routinely decomposed into large and small scale patterns (Holl, 1963) to
clarify the results of large-scale circulation changes, fluctuations in
upwelling, etc. The large-scale SST pattern derived from Figure 2 is show
in Figure 3. Its similarity to ocean current systems will be pointed out
later. i

B. SEA AND SWELL

The analysis and forecast program for wind waves uses a singular
technique to obtain significant wave height and period. Surface geostrophic
winds at 3-hourly intervals are the basic input. Duration is determined to
the nearest 3 hours and fetch corrections are made in regions of offshore flow.
The formulae for wave height and period as functions of duration Dp and
geostrophic wind speed Ug used at FNWF are

H1/3
H

2
a(Ug)” pp + bUg

1/3 (c +dDp) Ug te

A sample wave analysis is show in Figure 4.

Swell is defined as waves which have traveled more than 24 hours from
a generating area. Based on a history tape of wave heights, periods and
directions at 12-hourly intervals; travel distance, swell height and swell
period are computed from the following equations:

D= mn
ay T, mt

4
i
—4
i)
+
re
=)
Oo
>
Ni-

2h6

Sch. Lleies ANAL
‘ os 0 s
fe id PROJECTION: POLAR STEREOGRAPHIC—TRUE AT 60 NORTH LATITUDE . 2 FLEET NUMERICAL WEATHER FACILITY

SCALE: 1:60,000,000 or ef MON INTERES CALIFORNIA

Operational Sea Surface Temperature (SST) Analyede
for 00 GMT 1 February 1965. Degree C.

“Figure 2.

Q8Z 01 FEB SS

»6R

Gaerne i) "

SEA. SL ANAL. 22 21 EEB 65 -

Figure 3. Large Scale Part of SST Analysis for 00 CMT
1 February 1965. Result of Scale and Pattern
Separation,

2h8

WHAT, ANAL "62-01 FEB 6S
Figure 4. Wind Wave Analysis for 06
Heights in Feet.

A Falc

GMT 1 February 1965.

Significant

= 2h9
where D is travel distance, T, is the period at the end of fetch, m is the

mean map factor, t is decay time, T, is the swell period, H, the swell
height. H, the height at end of fetch, and a, and b; are “constants.

Swell analyses and forecasts are plotted in the same manner as the waves in
Figure 4.

OCEAN CURRENTS

The details of this program have been described earlier by the author
(Hubert, 1964). Essentially, the computational procedure accounts for two
principal current components -- (1) the "characteristic" or thermohaline
flow and (2) the mass transport due to wind and waves.

Assuming a level of zero current velocity at some depth (Az) , the
geostrophic thermal current at the surface is computed from the mean tempera-

ture, T, in the layer

EAT,

=> = —

pe St ey ee eee Oa
S f

In practice, the mean temperature is obtained from a weighted combination
of a climatological temperature field at 200 meters and the synoptic SST
Analysis described earlier.

The wind-driven current as determined by Witting (1909) is obtained
from

—=>
where Me is the mean geostrophic wind speed for a 36-hour period.

Figure 5 is an example of a current transport chart (in nautical miles
per day) obtained at FNWF on a synoptic basis. As can be seen from this
figure, well-known features such as the Gulf Stream, Kuroshio, Equatorial
Counter Current, etc., are quite well defined by this procedure. Since the
computations are carried out in component (u,v) form, directional fields are
also available.

In order to obtain a single continuous field displaying both direction
and speed of the computed currents, a stream funetion ()analysis is made
using methods similar to those employed by Bedient and Vederman (1964) to
represent atmospheric flow in the tropics. ‘he vorticity of the current flow
is determined from the (u,v) component fields and the Poisson equation

whe Oe. oe
ox oy

is solved for “W using relaxation techniques.

250

. s ‘ns 5 j oo EK fe Sena Le! en aN She a
. _ p ‘om ‘ SRN : me . 4 SUarEN TRANS OR : NM /DAY a ;
Gana i PR cies Pe a dee a}

Figure 5. Current Transport Computation at 06 GMT 1 February 1965.
Transport in Nautical Miles/Day. Transport Over 12 n.mi./day

Stippled.

251

The stream function field which corresponds to the current transport
chart in Figure 5 (06 GMT 1 February 1965) is show in Figure 6. Current
vectors have been plotted at selected gird points to show the degree of fit.
The fact that the derived stream function is nondivergent while there is
divergence in the initial velocity field explains some of the cross-contour
flow. In general, however, this appears to be small in most places, and the
stream field provides a good representation of the current pattern.

It is interesting to note that the stream function analysis shows close
correlation to the large-scale SST analysis shown in Figure 3. As one
should expect, thermohaline considerations (as influenced by the semipermanent
circulation of the atmosphere) determine the large-scale current pattern while
mass transport by wind and waves contributes toward smaller scale details.

A good example of the latter effect can be seen to the northeast of
Hawaii. A strong, quasi-stationary cyclone has completely disrupted the
normal west-east extension of the Japanese current. To a lesser degree, the
same thing is happening off the west coast of France; the typical northwesterly
current has been replaced by flow from the southeast.

D. ADVECTIVE TEMPERATURE CHANGES

From the computed currents one can determine the change in SST which
would be due to advection alone. Since the "permanent" or thermohaline
component would be nearly along the sea surface isotherms, the advective
patterns should result primarily from atmospheric driving forces of synoptic
scale. Figure 7a shows the results of advecting the SST pattern with the current
field show in Figures 5 and 6. Areas of cold advection are stippled with
heavier shading used to denote advection of greater than about O.1F per
24 hours. Weak warm advection is indicated by lack of shading while warm
advection stronger than 0.1F per 24 hours is shown by cross-hatching.

Figure 7b represents the actual difference between the SST analyses
from 12 GMT 1 February and 12 GMT 31 January 1965. As can be seen, the cor-
respondence in some parts (particularly in the vicinity of strong storms
near Hawaii, Newfoundland and the Azores) is fairly good. In other areas
the signs are clearly opposite. Ome can only conclude that advection plays
an important role in some areas and is completely outweighed by heat exchange
effects in other areas. Namaias (1959) and Eber (1961) came to about the
same conclusion. The maximum advection computed in this case was 0.85°Rr per
24 hours which agrees with earlier findings of Laevastu (1960).

5. THE FNWF SCHEME FOR SUBSURFACE THERMAL STRUCTURE ANALYSIS AND
PREDICTION

Figure 8 is a further breakdow of the master scheme discussed in
section 3. As can be seen, all of the computational programs described in
the preceding section enter into the determination of thermal structure with

252

Ned

FUNQTION™ .2

~. STREAM>

.
* +

Figure 6. Current Stream Function Analysis for 06 GMT 1 February 1965.
Units 101 sec~+. Current Arrows to Scale Plotted at

Selected Grid Points.

253

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SCHEME FOR SUBSURFACE
THEBAMAL STRUCTURE
ANALYSES ANID FORECASTING

SEA & SWELL
REPORTS

M.L.D. BY
WAVE MIXING

ATMOSPHERIC
ANALYSES

WINDS, TEMPERATURE
HUMIDITY, CLOUDS

CONVERGENCE |.
8N7en DIVER‘NCE

HEAT BUDGET
COMPUTATIONS

CONVECTIVE
STIAAING

SEA-SURFACE TEMPER-
ATURE ANALYSES (S57)

PREVIOUS ANALYSES, | SUBSURFACE THERMAL STRUCTURE
HVDROCLIME DATA, ANALYSES (M.L.D, THERMAL GRAD-
AND PT. REPORTS | ENT, TEMP AT STANDARD DEPTHS —

ATMOSPHERIC FORECASTS WAVE, CURRENTS,

} HEAT BUDGET, AND
EED BA 5.5.1 FORECASTS

Cee
MOSPHERIC FORECASTS
SUBSURFACE THERMAL |
STRUCTURE FORECASTS

pal

Figure 8. Fleet Numerical Weather Facility, Monterey Scheme for Subsurface
Thermal Structure Analyses and Forecasting.

ay)

depth. The analyses will be built dowmward from the surface (where the
most reports are available).

The previous day's analysis will first be modified by mechanical and
convective mixing, where applicable, and large-scale convergence and
divergence in the surface layers. Climatological (hydroclime) restraints
will be used to keep computed changes within reasonable limits. Finally, BT
observations will be introduced by means of a median seeking technique such as
used in the SST analysis program.

Most of the components needed to assemble the complete subsurface
analysis and prediction package have been programmed for numerical solution.
The first major portion to be completed is expected to be a hemispheric
analysis and prediction of mixed-layer depth.

256

Bedient, H.A.
Vederman, J.

Eber, L.E.

Holl, M.M.

Hubert, W.E.

Laevastu, T.

Namias, J.

Witting, R.

Wolff, P.M.

and

REFERENCES

1961:

1963:

1964:

1960:

1965:

1959:

1909:

1964:

Computer analysis and forecasting
in the tropics. Monthly Weather
Review, Vol. 92, No. 12, pp 565-577.

Effects of wind-induced advection
on sea surface temperature. J.
Geophys. Res., 66, pp 839-844.

Scale-and-pattermn spectra and
decompositions, Tech. Memo No. 3,
Meteorology International, Monterey,
Calif.

Computer produced synoptic analyses
of surface currents and their ap-
plication for navigation. Presented
1964 National Marine Navigation
Meeting, San Francisco, Dec 1964.

Factors affecting the temperature
of the surface layer of the sea.
Merentutkimuslaitoksen Julkaisu,
167, pp 131.

Synoptic scale heat exchange and
its relation to weather. FNWF,
Monterey Tech Note No. 7, Feb 1965.

Recent seasonal interactions between
North Pacific waters and the over-
lying atmospheric circulation, J.

Geophys. Res., 61, pp 631-646.

Zur Kenntriss des vom Winde erzengten
Oberflackenstromes. Ann. Hydrogr.
Marit. Met, 73:193.

Operational analyses and forecasting
of ocean temperature structure.

(Rpt Fleet Numerical Weather
Facility)

257

SYNOPTIC SCALE HEAT EXCHANGE AND ITS RELATIONS
TO WEATHER

Taivo Laevastu
U. S. Fleet Numerical Weather Facility
Monterey, California

"The opinions presented in this paper are those of
the author and do not necessarily represent the
official views of the Navy Department at large."

258
ABSTRACT

The present study was undertaken to learn about the contribution to
heat exchange contrasts through sea-air heat exchange and to contribute
therewith to the understanding of the feedback of energy between the sea and
the atmosphere on a synoptic scale.

The formulas and procedures for computation and forecasting of the
heat exchange components are given and reference is made to the studies of
their accuracy and sources of errors. The disadvantages of the monthly and
seasonal computations, as compared to synoptic ones, are briefly discussed.

Examples of the synoptic distribution of heat exchange components
during given days over the North Pacific Ocean are presented. Graphical and
descriptive models of the heat exchange patterns in relation to anticyclones
and different developmental stages of cyclones are constructed. Based on
these models, the effects of the energy change on the ocean surface properties
are explained and verification demonstrated with synoptic analyses of the
short-term changes and anomalies of sea surface temperature. The return
effects of these anomalies to the surface weather are postulated and the
numerical tests of the use of "correction factors," indirectly derived from
the present study, are briefly indicated.

A hypothetical model of the coupling of the heat exchange model with
the 500-mb SD patterms is given, its use for deriving surface pressure pat-
terns is demonstrated, and the possible use of this approach is shown by
verification of forecasting attempts over 48 and 72 hours.

Though some principal aspects of the presented feedback models have
been tested and found to contribute toward improving the present forecasting
models, they are still experimental in nature. However, they increase the
prospects of preparing numerical 3 to 5 day forecasts in.the not too distant
future. The primary use of heat exchange computations at Fleet Numerical
Weather Facility is in synoptic oceanographic analysis and forecasting.

259
INTRODUCTION

The development of truly synoptic oceanographic analysis and forecast-
ing emphasizes thermal structure in the sea and therefore requires the
quantitative knowledge of energy exchanged between the sea and the atmosphere.
It can be postulated that temperature in oceanographic analyses has the same
importance as atmospheric pressure in meteorological analyses. Furthermore,
oceanographic analysis and forecasting must be based mainly on synoptic weather
observations by ships, as truly synoptic subsurface oceanographic observations
are scarce indeed and would be too time consuming and expensive to make on a
worldwide synoptic scale.

A number of meteorologists, especially those from the sorcalled Bergen
School and a few others, have left no doubt that there is also a need to include
heat exchange effects into successful weather forecasting models. Therefore,
the synoptic study of heat exchange finds application also in meteorology.

A number of heat exchange studies have been done in the past on a
seasonal and monthly basis; however, synoptic studies have been scarce. The
most extensive of the latter are by Petterssen, Bradbury and Pedersen (1962)
and by the present author (Laevastu, 1963).

Among the objectives of the present study, reported herein, were:

(1) To investigate the feasibility of synoptic computations of heat
exchange components and to study their accuracy and possible sources of
errors.

(2) To study the day to day variability of heat exchange patterns
at the surface and their relations to surface Weather and upper air patterns.

(3) To study the effects of heat exchange on the ocean. This report
presents mainly the results of objective (2) above; other objectives are
dealt with briefly. In addition, a hypothetical model of energy exchange
and feedback between the ocean and atmosphere is given, and its possible
application is explored. This study was carried out with the support of NSF
Grants Nos. GP-353 and GP-2459 and with support from Fleet Numerical Weather
Facility. The author wishes to express his sincere thanks to Commander
Hubert and Mr. Carstensen for valuable advice and help in the preparation of
this paper.

FORMULAS FOR HEAT EXCHANGE COMPUTATION AND
ACCURACY AND SOURCES OF ERRORS

The formulas used for heat exchange computations are summarized in
Table 1. The validity and accuracy of these formulas have been described
earlier by the present author (Laevastu, 1960). The heat exchange

260

(1)
(2)
(3)
(4)

(5)

(6)

TABLE, la.
Formulas for synoptic computation of heat exchange
between the sea and the atmosphere
All units of in g cal em-> oy hot

Insolation ( Q5 ) = O0.O14A,tg (1 - 0.006 c3)
Albedo (Q,.) = 0-15 Qs - (0.01 Q,)®
Effective back radiation (Q) = (297 - 1.861, - 0.95U,)(1-0.0765 C)
Latent heat transfer

@,, “25 POS- Q = (0.26 +0.077 V)(0.98 e,-e,

Gy “fg neg. Q, = 0.077 V (0.98 e,, -e,)
Sensible heat transfer

T, -T, Pos. Q, = 39 (0.26+0.77 V)(Z, - 7.)

aaa ogo Gy SS SVC, S we.)

Total heat exchange (Q,) Fey > Cy oie Gs,

261

TABLE 1b.
List of notations used in Table la

noon altitude of the sun (° )-

cloud cover (in 1/10 of the sky)

water vapor pressure of the air (mb)
saturation vapor pressure of the sea surface (mb)
latent heat of vaporization (g cal ean)
effective back radiation / g cal (2un) "7
latent heat transfer

sensible heat transfer

reflected heat (albedo)

insolation

length of the daylight (min)

air temperature (°c)

sea surface temperature (°C)

relative humidity (%)

wind speed (m gee)

262

computations, which form the basis for the. present study, were made manually.
The computational procedure has been described in technical reports (Laevastu,
1963), where also nomographs are given. Before averaging the meteorological
elements (reported by observing ships) by areas (e.g., 2 1/2° or 5° squares),
a subjective contouring of the distributions is necessary. This subjectivity
will be eliminated in computer programs of synoptic heat exchange now in
preparation at Fleet Numerical Weather Facility, Monterey.

Detailed discussions of the sources and magnitudes of errors are given
in the aforementioned technical reports. It has been concluded that the
plausibility of the results obtained depends largely on the density of
meteorological elements reported over the ocean and on their analysis because
not the absolute values of the elements, but rather the differences between
the properties at the sea surface and some higher observation level (e.g.,

8 to 10 meters), are used in most cases. Other difficulties, such as the
analysis of distribution of cloud cover, estimation of wind speed over the
sea, etc., are well know.

In the evaluation and use of the computed results one must also be
conscious of the possible effects of short-term (diurnal and interdiurnal)
variability of meteorological elements. However, as will be shown later, the

patterns of heat exchange components are usually larger in scale and relatively

persistent from day to day, changing in intensity and position in relation to
the change of corresponding surface pressure patterns, thus eliminating partly
the effect of short-term variations. The effects of diurnal fluctuations of
meteorological elements on computed heat exchange are discussed by the present
author elsewhere (Laevastu, 1960, 1963).

Considering that rather large variability in meteorological conditions
occurs from day to day, that many of the heat exchange formulas use the dif-
ferences between the sea surface and the surface air properties, that the
relations of the influence of the elements in these formulas are often non-
linear, and that the formulas have mostly been worked out for and verified on
short-period (e.g., 24 hours) measurements and computations, it becomes
obvious that synoptic computations of heat exchange are usually not comparable
to heat exchange patterns computed with monthly or seasonal averages. The
above described condition is numerically illustrated in Table 2. Four values
of sea-air temperature and water vapor pressure are arbitrarily selected.
Furthermore, four values of wind speeds are taken in two different sequences,
which give two different values of sea-air exchange (Q ) in the last two
colums. The average values of meteorological elements as well as the
averages of the computed sea-air exchange are given on, the sixth line. The
averaged Qa, walues are 59.3 and 182.5 g cal cm ~ 24h ~ and illustrate the
differences in results caused by a difference in the sequence of wind speeds
used, even though the average wind speed is 6.25 m sec ~ in both cases. The
seventh line gives the Q, value (104) computed with the averaged values of
meteorological elements. It can be concluded from Table 2 and from a
knowledge of the variability of meteorological conditions that monthly and
seasonal computations can be expected to yield only an approximate distri-
bution of heat exchange patterns and their relative numerical values. To

263

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264

arrive at realistic monthly or seasonal average heat exchange values the
synoptic heat exchange must be computed with synoptic meteorological elements
and summed and averaged after these computations have been performed.

EXAMPLES OF SYNOPTIC DISTRIBUTION OF HEAT EXCHANGE COMPONENTS

Some examples of the distributions of heat exchange components are
shown in Figures 1 to 7. The variations in insolation are largely determined
py the time of year, latitude and cloud cover (Figure 1). The effective back
radiation is also affected by cloudiness, the relative humidity and the sea
surface temperature (Figure 2). The transfer of latent heat and sensible heat
transfer are most directly connected to surface weather systems (Figures 3
and 4). These two components are summed together and the quantity called sea-
air exchange Con ). Two consecutive days of Q, distribution over the North
Pacific are shown in Figures 6 and 7 together with the surface weather charts
and the total heat exchange (Qp ). These figures indicate that the patterns
of Qa are large in scale, corresponding to surface weather patterns, and that
they change from day to day in the same manner as the weather patterns. In
lower latitudes, below 20° North, a number of smaller centers of Qa are
apparent. This is partly in agreement with the surface pressure distribution
as shown by synoptic analyses. Furthermore, there are some uncertainties
involved in the computations in this area, as the ship reports south of about
20°N, as well as north of about 55°N, are sparse. The higher values of Q,
are usually found on the cold air side of cyclones. Low or negative values of
Qa are found, on the other hand, in the warm sectors of cyclones. Relatively
high Qg gradients can be found along the cold fronts. The relations of the
Q, fields to cyclones change slightly with the age of the cyclone itself.

In some cases pronounced Q, contrasts have been observed in areas and
places where cyclogenesis would be expected and does, in fact, occur 1 to 2
days later. In the eastern parts of the ocean the Q, contrasts are much less
pronounced than in the western parts. The disappearance of Q,_ contrast within
& cyclone area usually precedes the dissipation of the cyclone, and might be
useful in forecasting the filling of a system. Physically, it can be explained
that the energy sources and sinks for the cyclones at the sea surface are cut
off due to the exchange processes at the surface. These tend to diminish the
gradients of properties between the sea surface and the lower layers of the
atmosphere and consequently to dissipate the cyclone because it has been cut
off from any new sources of energy.

The relations between anticyclones and heat exchange patterns are less
distinct than in the case of cyclones. However, there is usually a higher
Q, in the eastern part of an anticyclone and lower Qq pattern in the western
part of it. Most of the well defined high and/or low Q_ patterns are related
both to the cyclone and the adjacent anticyclone (see further Figure 9).

265

Received radiation.

(Insolation minus reflected
radiation, Q,—Q, )

9 cal cm* (24h)-!
May 14, 1957

Figure 1. Received radiation on 14 May 1957

ee = “es Effective back radiation. Qp

g cal cm (24h)-!

May 14, 1957

|
<|Q0

(Sind 6

Figure 2.

Effective back radiation on 14 May 1957

SS Exchange of sensible heat. Qp

(Positive values indicate that
the air gains the heat and
Negative values indicate
heat gain by the sea.)

g cal cm (24h)-!

e ce

cba

Figure 3.

Transfer of sensible heat on 14 May 1957

266

the sea surface.)

g@ cal cm? (24h)7!

May 14,1957

Figure 4. Transfer of latent heat on 14 May 1957

Total heat exchange. Q,
(Positive values indicate that the
sea gains and negative values
indicate that the sea loses the heat.)

g cal em224h)'
Feb. 14, 1957

300s

Le |
Lea

Total heat exchange. Q,
(Positive values indicate that the
sea gains and negative values
indicate that the sea loses the heat.)

gcal em&{24h)'

Aug. 14,1956

267

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LO <I ZZ)

=p
Ke |

Total heat exchange. Q,
(Positive values indicate that the

Ma Ao
we

Figure 6. Surface pressure, sea-air exchange and total heat exchange
on 16 May 1957.

268

Synoptic Weather Map
Sea Level, |1230GMT

May 17, 1957

Td TOP a D ~
Heat exchange between sea and

atmosphere over North Pacific.
Air-sea-heat exchange. (Q)+ Q +Q,)
(Positive values indicate that the air gains

the heat and the negative values indicate

heat gained by the sea)

9 cal cm4(24h)!

May 17, 1957

Total heat exchange. Q,
(Positive values indicate that the
seq gains and negative values
indicate that the sea loses the heat.)

9g cal cmrA{24hr'!
May I7, 957

Figure 7. Surface pressure, sea-air exchange and total heat
exchange on 17 May 1957.

269

The total heat exchange Qp indicates the amount of heat gained or lost
by the sea in 24 hours. Figure 5 shows the total heat exchange distribution
at 2 given days in the winter and summer, respectively. High loss of heat
along the east coast of Asia can be observed during the winter over an area
which reaches to a relatively low latitude (ca. Doon) and a considerable
distance out from the coast (to ca. 165°E). Eastern and southeastern parts
of the ocean show relatively low loss or gain of heat during the same date.

The distribution of Qp in summer is more latitudinal and is affected by
meteorological conditions as shown by Figures 5, 6, and 7.

SIMPLIFIED PHYSICAL MODEL RELATING HEAT EXCHANGE PATTERNS TO
CYCLONES AND ANTICYCLONES

Based on the examination of a number of synoptic heat exchange computa-
tions over the North Pacific a simplified physical model, relating the heat
exchange patterns to cyclones and anticyclones, is presented in Figure See this
model is largely self-explanatory. It shows the location of the heat exchange
patterns with the developmental stages of the cyclone as well as in the vicinity
of the surrounding anticyclones. The adjacent cyclones and anticyclones are
connected to each other through "common" heat exchange patterns, as schematically
shown in Figure 9. Figure 10 presents a hypothetical vertical E-W section of
a cyclone as it might be related to the heat exchange processes. Figure 11 gives
a scheme of heat exchange, flow patterns and upper level topography of a warm
sector cyclone. Whether the simplified models in Figures 8 to 11, derived from
an examination of 30 synoptic Qg charts, correspond exactly to the processes in
nature or need further testing and improvement is not subject to discussion here;
their importance lies in the fact that they fit without serious discrepancies
and lead to the consideration that the sea-air heat exchange patterns might be
related to upper air patterns. In fact, a study of the 500-mb small scale (SD)
patterns shows a remarkable similarity to the heat exchange patterns, especially
with regard to position. Tests were made, therefore, using the 500-mb SD pat-
terns for analysis as well as forecasting of surface pressure distribution.

(The pattern separation procedure has been described by Holl, 1963). The con-
struction of surface flow patterns from 500-mb SD patterns is shown in Figure
12. The positions of surface lows and highs and their central pressures are
also indicated, using preliminary relations between SD central values and
central pressures at the surface (Figure 13). The results of another preliminary
study show that cyclone centers should be sloped about 4° latitude ENE of SD
centers and anticyclone centers about 5° toward 120°. As the 500-mb forecasts
usually show somewhat better skill over longer forecasting periods (48 to 72
hours), this model has led to a possible auxiliary surface forecasting procedure.
Some numerical results of verification and comparison tests of this model for

48 and 72 hours compared with the operational Fleet Numerical Weather Facility
surface model are shown in Table 3. A report on the incorporation of the above
results in extended numerical forecasts (3 days) is in preparation at Fleet
Numerical Weather Facility.

270

High positive Q.,

3
area of pressure rise area of pressure fall

Frontal

High Q_
a

Low or negative Q,:

Low or
negative Q>
pressure
decreasing

High positive Q,:

pressure rising

wave Warm sector cyclone

Low Q.,
pressute fall

Hig Q
pressure rise

Fully occluded cyclone Anticyclone

Figure 8.

Simplified model on the relations between sea-air heat
exchange and cyclones and anticyclones

Low or
negative

Schematic

Figure 9. Scheme of Qo patterns common in adjacent cyclones and
anticyclones.

ene

"Inflow" and convergence
on'ground. "Outflow" aloft

Release of latent heat
ieee heating of clouds

= Flow

os — Isothermal
————
surface

= _

~-_—-——

Q, Weel ative
Zia ll TTVTTr

Q_ positive Heating and uptake
a of water vapor

Figure 10.

High

positive Q,

Upper level
(e.g. 500 mb) pattern

Negative Q,

Upper level flow

Figure ll.

Figure 10.

Figure 11.

Schematic Vertical E-W Section Through a Warm Sector
Cyclone With Assumed Model on the Relation to Heat Exchange.

Scheme of Heat Exchange, Flow Patterns and Upper Level
Topography of a Warm Sector Cyclone.

273

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Central pressure (F), mb

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and central pressures of cyclones and anticyclones.

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6
SIMPLIFIED FEEDBACK BETWEEN THE SEA AND THE ATMOSPHERE

The ocean/atmosphere feedback has been described in general terms in a
few publications in the past (e.g., Bjerknes, 1960), but a complete descrip-
tion in terms of energy exchange is not available. In this chapter an attempt
is made to describe the synoptic feedback between the sea and the atmosphere,
based upon the results of previous chapters.

An essentially diabatic simplified model is considered herewith, with
the conversion of energy in the atmosphere itself being largely neglected,
as it has been described by several authors earlier. Some existing numerical
meteorological models permit heat exchange input without extensive changes.
Tests with these models will clarify further the quantitative relations between
energy exchange and weather processes and will indicate the usefulness of the
inclusion of heat exchange parameters.

The feedback system can easily be derived from Figures 8 to 11, a simpli-
fied description of which is given below. In this description it is assumed
that the sea surface isotherms are essentially latitudinal and equidistant at
the initial stage. Compass directions, rather than descriptive terms of the
sectors of cyclones and anticyclones, are used.

Anticyclonic circulation:

Atmospheric circulation in relation to heat exchange. In the N part of
the anticyclone the eastward flow is approximately parallel to the sea surface
isotherms and would not result in appreciable sea-air exchange. In the NE
part of the anticyclone the air flow has a southerly component, which brings
colder air over warmer water. This factor can increase progressively in the
E and SE parts of the anticyclone. A slight pressure rise is usually observ-
able in the SE part of the anticyclone in the following days which gives an
apparent movement of the anticyclone to the SE, and an apparent location of
anticyclone center towards areas with negative sea-surface temperature
anomalies.

The energy gain of the air in S and SW part of the anticyclone decreases,
and low or negative Q, values are found in the NW sector due to the north-
ward component of the air across sea surface isotherms. In this area pres-
sure fall occurs.

Changes in the ocean, caused by advection and heat exchange. If the heat
exchange and advectional effects in the sea and in the atmosphere were perfectly
balanced, few disturbances would result. However, due to local inequalities
of heat exchange processes and to advectional changes sea surface temperature
anomalies are caused, which in turn may affect the heat exchange processes.

In some cases advectional effects might account for the greater part of
the anomalies in the oceans. In the W and NW part of the anticyclone, warmer
weather is advected towards N and NE. In the NE and E part colder water is
advected towards south. Further cooling of the ocean in the SE sector can
be effected by intensive sea-air exchange in this area. In the S and SW part

; 2//
of the anticyclone, slight warming of the sea surface is expected partly due
to latitudinal warm advection and partly due to low sea-air exchange and
high insolation.

Subsequent changes in the atmosphere. Assuming that the advectional and
local heating effects, described above, are not in equilibrium with the cool-
ing and heating processes and advectional anomalies result, the following
further changes could be expected in the atmosphere:

In the W and NW part of an anticyclone the warm advection would counter-
act the heat loss of the air by decreasing sea-air temperature difference,
and the expected pressure fall would be slowed down with time. If the warm
water advection is especially strong, a pressure rise in the N part of the
anticyclone could be observed in a few days, resulting in an apparent slight
northward movement of the high.

In the NE and E part, the cold water advection would diminish and/or
counteract the increase of sea-air exchange. The same effect occurs in the
E and SE part, thus counteracting the prospective accompanying pressure rise
and eastward movement of the high. This mechanism might explain the quasi-
stationary nature of highs in the lower latitudes over the oceans.

Cyclonic circulation:

Atmospheric circulation in relation to the heat exchange. In the W and
SW part of a cyclone, high positive sea-air exchange is taking place, due to
the southward component of cool, drier air across the sea surface isotherms.
This cool air is usually accompanied by subsidence which results in clear,
cloudless skies. The high sea-air exchange is accompanied by relatively rapid
pressure rises in this sector.

In the S part of a cyclone, the air flow is nearly parallel to the sea
surface isotherms or with a slight northerly component, which increases in
the SE part of the cyclone and results in heat and moisture loss by the air,
accompanied by pressure falls.

In the E and NE part the flow is towards colder sea surface temperatures.
The heat loss by air is, however, decreasing rapidly, because of cooling of
the lower layers of the air and creation of stable conditions.

In the N part of the cyclone the curvature is usually relatively sharp
and in the NW part the heating of the air starts again due to a slight
southerly component.

Changes in the ocean caused by advection and heat exchange. In the W
and SW part of a cyclone, cool sea surface advection takes place. This cool-
ing is further aggravated by deep mixing in the sea due to heavy winds and
waves and by heat loss from the sea through sea-air exchange.

In the S and SE part an E to NE advection of warmer water occurs, which
is accompanied by low heat loss or occasionally heat gain through sea-air

278

exchange (diminished by low insolation due to heavy cloudiness behind a warm
front).

In the NE and N part this advection turms to cold advection from the
NE. Slight upwelling can be expected in the center of a cyclone due to
divergence in the surface layers.

Subsequent changes in the atmosphere. In the W and SW part of the

cyclone, the cool advection counteracts the normally large sea-air exchange
and warming of the air, which results in further SE movement of the cold front.

oe Ay A A

In the S and SE part the warm advection woudid extend the area of heat loss
toward the NE, accompanied by a pressure fall in the same direction. It is
thus apparent that the advective and local changes cause a kind of twisting
movement around a cyclone giving a force in the NE direction. The model of
occluded cyclones together with the description above also gives partial
explanation why occluded cyclones tend to move more to the left. The above
model also explains partly the tendency of cyclone centers (and the W parts
of them) to locate over areas with positive sea surface temperature anomalies.

This description of the feedback model should not be considered as a
final one, but rather as a framework, subject to further refinement. The
observed pressure tendency anomalies, empirically related to sea-air exchange,
remain unexplained in detail in these models.

The above described model should be thought of as entirely auxiliary to
existing operational models of atmospheric behavior, providing an additional
"correctional" factor only.

Partial verification of the above described feedback model has been
sought and indeed found. The effects on the ocean are demonstrated in
Figures 14 to 16. Figures 14 and 15 show surface pressure distribution and
the positions of sea surface isotherms on 10 and 13 December 1964. Figure 16
shows the sea surface temperature changes at four selected grid points in this
area and period, taken from Fleet Numerical Weather Facility's synoptic
analysis of sea surface temperature ( Wolff, 1964).,. Positions of these grid
points are indicated on Figure 14. An examination of Figures 13 to 15 indicates
that the observed short-term sea surface temperature changes are partly
advectional and correspond to the changes described and expected in the above
simplified feedback model. The partially advective nature of the short-term
sea surface temperature changes is furthermore substantiated by the synoptic
analyses of surface currents (Hubert, 1964). Furthermore the change of sea
surface isotherms between 12 and 14 December 1964 (Figure 17) also indicates
that the sea surface temperature change patterns are large in scale and
correspond to expected advectional patterms, as show by a comparison to surface
pressure analysis charts for this period.

The preliminary verification tests of the atmospheric part of the feed-
back model, utilizing the short-term sea surface temperature anomalies, was |
done subjectively for a number of forecasting periods. The results indicated

279

Figure 14. Surface pressure and sea surface isotherms in the NT Atlantic on
10 December 00Z 1964. The subsequent movement of cyclone and
anticyclone centers is shown with arrows.

Figure 15. Surface pressure in the NW Atlantic on 13 December 00Z 1961.

280

59
58
37
10 XII
ty
{o)
e
2
2 o7
oO
=
2 66
65
64
63
62
10 XII
Figure 16.

12 14 XII

{

12 14XII

Date

65
64
63

70
69
68

67
66
65

Change of sea surface temperature at four selected
grid points in the NW Atlantic from 10 to 14 December

1964.
Figure 14.)

(Positions of the selected grid points see

281

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282

some skill. Consequently some slightly different "correction factors were
programmed at Fleet Numerical Weather Facility for numerical testing in
connection with operational numerical surface forecasting models. These tests
and developments, promising minor improvement to existing models, are in
progress and will be reported at a later date.

Although the synoptic energy feedback from the sea to atmosphere promises
some use and improvement, the principal use of synoptic energy exchange
computations at Fleet Numerical Weather Facility is in oceanographic analysis
and forecasting, as demonstrated by Commander Hubert at this conference

(Humbert, 1965).

Bjerknes, J.

Holl, M.

Hubert, W.E.

Laevastu, T.

Petterssen, S.5; D. L. Bradbury
and K. Pedersen

Wolff, P.M.

283

REFERENCES

1960

1963

1965

1964.

1960

1963

1962

1965

1964

Ocean temperatures and atmospheric
circulation. WMO Bull.
NS) eisilai/c

Scale-and-pattern spectra and
decomposition. Met. Int. Inc.,
Monterey, Techn. Memo 3.

U.S. Fleet Numerical Weather
Facility activities relating to
sea-air interactions on a synoptic
scale. FNWF Techn. Note No. 5.

Computer produced synoptic analyses
of surface currents and their
application for navigation.
Presented at the 1964 National
Marine Navigation Meeting, Institute
of Navigation, Dec. 7-8, 1964,

San Francisco.

Factors affecting the temperature
of the surface layer of the sea.
Soc. Scient. Fennica, Comment.

Physico-Mathem. 25(1):1-136-

Energy exchange in the North
Pacific; its relations to weather

_ and its oceanographic consequences.

Parts I, II and III. Hawaii Inst.
Geophys. Rpts Nos. 29, 30 and 31
(Rev.).

The Norwegian cyclone models in
relation to heat and cold sources.

Geofys. Publ. (Geophys. Norv.
24(9):249-250. ;

U.S. Fleet Numerical Weather Facility;
The use of 500-mb SD pattern fore-
casts for derivation of surface
forecasts, and the accuracy of 3-day
forecast based on 500-mb SD patterns.
FNWF Techn. Note(in preparation).

Operational analyses and forecasting
of ocean temperature structure. Ist

U.S. Navy S . on Military Oceanogr.,
ErOCrs TYI-1o5

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LABORATORY STUDIES OF WIND ACTION ON WATER STANDING IN A CHANNEL

G. M. Hidy
National Center for Atmospheric Research
Boulder, Colorado

and
E. J. Plate

Colorado State University
Fort Collins, Colorado

285

286

ABSTRACT

The processes of wave and current development resulting from wind action
on initially standing water have been investigated in a wind-water tunnel.
The mean air flow over wavy water was examined along with the variation of
several properties of the water motion with fetch, water depth, and wind
speed. Measurements of phase speed and length of significant waves, the
standard deviation of the water surface, the average surface drift, the auto-
correlation of surface displacement and the frequency spectra are reported.
The experimental results indicate that (a) the air motion in the channel
follows a three dimensional pattern characteristic of wind tunnels of rectangular
cross-section; (b) wind waves generated in the channel travel downstream at
approximately the same speed as gravity waves of small amplitude, provided the
effect of the drift current is taken into account; (c) the average drag co-
efficients for the action of the wind on the water surface increase with in-
creasing wind speed, and these data are reasonably consistent with results of
previous investigators; (d) the autocorrelations and frequency spectra indicate
that the wind waves in the channel consist of nearly regular primary waves on
which are superimposed smaller ripples; (e) energy in the high frequency range
in the spectra tends to approach an equilibrium distribution while the lower
frequency components continue to grow with increasing fetch; and (f) a
similarity shape for the frequency spectra develops. The experiments in this
study were not intended to model the processes of interaction between the
ocean and the atmosphere. Nevertheless, the small waves generated in the chan-
nel appear to be at least qualitatively related to the development of waves
on much larger bodies of water.

287

I. INTRODUCTION

In spite of a long history of effort devoted to the air-water interaction
problem, the basic knowledge of the mechanisms for transport processes near
the boundary between these two fluids has developed rather slowly. A variety
of theoretical and experimental studies have been reported in the literature,
but, because of the complexities of the physical processes involved, the
detailed nature of the interaction remains inadequately understood.

Most of the experimental studies of air-water interaction have been
undertaken on lakes or on the ocean where the conditions of the fluids are
highly variable in time and space. These investigations have contributed
significantly to the knowledge of the atmosphere and the sea. However, their
usefulness in elucidating the fundamental physics of the exchange processes
occurring between the two fluids is limited. Therefore, more studies should
be carried out under controlled conditions in the laboratory to gain new
insights into the mechanisms of transport across the air-water boundary.

Ursell (1956) has reviewed the fundamental laboratory experiments dealing
with air-water interaction that were undertaken before 1954. Since the
publication of Ursell's paper, a number of new investigations have been
reported which included those of Sibul (1955), Cox (1958), Fitzgerald (1963),
Schooley (1963) and Hanratty and coworkers (e.g., Cohen and Hanratty (1965)).
With the exception of Cohen and Hanratty's work, the experiments performed
by these investigators were not designed specifically to verify recent
theoretical conclusions, or to serve as a starting point for developing
refined ideas about the nature of air-water interaction. With this background
in mind, a detailed experimental program has been initiated at NCAR and at
CSU to study the relationship between the turbulent flow of air and water
in a channel.

Properties of the Fluid Motion

When air moves at moderate velocities over water, a drift current develops,
and small waves are generated on the liquid surface. A schematic picture of
the development of combined air and water motion along with the growth of
waves in a channel is shown in Figure 1. The properties of fluid motion
examined in this study are indicated in this drawing. The coordinate system
is indicated so that x is the distance downstream, and z is the vertical
direction. The mean water surface is given by z = d while the surface dis-
placement from this level is denoted as —& . The fetch F denotes the distance
from the leading edge of the water to a particular point somewhere downstream.
In terms of a two dimensional model, the velocity distribution in the water
is u(z), and the drift as the water surface is Up. The air flow is given by
U(z'), where Uso denotes the air velocity at approximately 20 cm above the
mean water surface, and z' = (z - d). The wave length } and the phase
speed ¢ denote properties of significant waves.

For the purpose of this study, significant waves will refer to the

288

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289

larger regular waves observed at a given fetch. In general, smaller ripples are
superimposed on the larger disturbances.

In this paper, a number of experimental results are discussed which
refer to the mean air and water motion as indicated in Figure 1. Measure-
ments of the statistical properties of the wind generated waves, including
the autocorrelation and spectral density functions, are examined in the light
of other properties of the fluid motion. There has been no attempt to model
the ocean-atmosphere interaction with the laboratory equipment. However, it
will be seen that a number of experiments for fluid flow in the channel are
at least qualitatively related to the observed small scale interaction between
the sea and the atmosphere.

II. EXPERIMENTAL EQUIPMENT AND PROCEDURE

The experiments were conducted in the Wind-water tunnel at Colorado
State University. This facility, shown schematically in Figure 2 consists
of a tunnel or a closed channel 0.61 m wide by 0.76 m high whose plexiglass
test section has a length of about 12 m. During operation, the maximum depth
of water is approximately 15 cm. Air is sucked through the tunnel at
velocities up to 18 mps by a large axial fan at the outlet. The inlet cone
is designed to give a y/, contraction ratio. Two fine mesh screens are
placed in the inlet cone. Honeycombs are placed just upstream of the outlet
diffuser to minimize the axial rotation in the air induced by the fan.
Sloping beaches are placed at the inlet and the outlet to prevent the reflec-
tion of waves. The "beaches" are constructed of aluminum honeycomb. The
inclines are shaped in such a way that as smooth as possible a transition can
be effected in the air-water flow. In this study, the bottom of the tunnel
was smooth.

The air flow through the tunnel was measured by a pitot-static tube
placed on a carriage in conjunction with a capacitative pressure transducer.
The probe could be positioned anywhere in the section of the tunnel from the
bottom to a level about 10 cm from the top.

The pressure gradient of the air and the depth of the water were measured
every 4 feet down the tunnel with piezometer taps connected to a set of
manometers .

Phase speeds and lengths of waves were determined from photographs taken
with a movie camera. The length for successive waves was measured from the
movies by comparing the distance between crests with a ruler in the picture.
The phase velocities of waves referred to a fixed point were estimated by
measuring from adjacent frames the distance traveled by a given crest during
the time between successive frames. Time intervals between frames were read
off a timer that was shown in the filn.

To measure the change in the height of the water, a capacitance probe
was used which is similar to Tucker and Charnock's (1955). This probe con-
sisted of a 34 gauge magnet wire stretched vertically along the center

290

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(0077 GY 01299 &
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<a
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t
——
Saas
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Figure 2. A general view of the wind-water tunnel.

291

line of the cross-section of the tunnel. These wires were placed at 1.2 m
intervals downstream from the inlet of the tunnel. The wire itself and the
water serve as the two plates of a condenser, and the insulation material
(Nyclad) on the wire provides the dielectric medium. The capacitance between
the wire and the water was measured with an AC excited bridge; the unbalance
voltage from the bridge was linearized, amplified and rectified so that a DC
output voltage was obtained which was directly proportional to the water
depth. The output signal was fed to an oscillograph where the gauge response
was continuously recorded during a run. The capacitance bridge-oscillograph
combination was calibrated to give a recorded amplitude linearly proportional
to the (varying) water depth with a flat response to frequency (+13) up
to approximately 30 eps.

From the continuous records of the surface displacement, data were read
off at equal intervals of 0.025 sec. These data were used for obtaining
values of standard deviation o of the surface displacement, the autocor-
relations R(t) of the surface displacement, and the spectral density function
o(f). The computations were carried out on the NCAR-CDC 3600 computer.

It was not possible to obtain the vertical velocity distribution in the
water. However, the surface velocity of the water u, was measured by placing
a small slightly buoyant particle on the water and observing the time required
for it to move past fixed stations downstream. Values of the surface velocity
could then be calculated from the intervals of distance of travel and the
time of passage.

In this study, attention was centered on the measurement of the properties

of water waves under conditions of steady (mean) air motion. In order to

_ attain steady conditions in the air flow, the wave development, and the set
up of water in the tunnel, the fan was started about 15-20 minutes before

the photographs, the pitot tube measurements, and the wave amplitude data

were taken at a particular location in the tunnel. In cases where wave data
were being measured, a sample of a wave train corresponding to the passage

of 100-200 waves was taken for a given run.

Samples of wave development were taken for several different conditions.
For the condition of water initially standing on a smooth bottom, air velo-
cities taken 20 em above the water surface, were varied from O to 17 mps,
and the depth of water was changed from 2.5 to 10 cm. The properties of
fluid motion in these cases were observed at distances of approximately 1.8
m to 12 m from the leading edge of the water.

III. THE AIR FLOW OVER THE WATER

Since the air is forced by the fan through the wind tunnel of approxi-
mately constant cross section, a pressure gradient develops in the down-
stream direction. The pressure in the air P, was found to vary approximately
linearly with fetch through the channel. Typical values of the pressure
gradient 1 oP, (cm water per cm) as measured in the last 6 m of the

pwg ox

292
channel are shown in Figure 3. The pressure gradient was found to increase

with wind speed, and with depth of the water.
Velocity Distribution in the Air

Measurements of the mean horizontal air motion in the vertical direction
and across the channel were taken at several sections for U,, from 6 mps to
about 14 mps. Typical data for vertical profiles along the center section
of the channel are show in Figure 4A. The vertical profiles of U(z')
indicate that the air flow generally develops a behavior characteristic of
turbulent flow in a boundary layer over roughened surfaces. In a few cases,

' a small kink in the distribution of U(z') was observed which usually appeared
at v5 cm height above the mean water level. Using pitot tube measurements,
Francis (1951) also observed these kinks. Schooley (1963) was able to find
the kinks by tracing mean trajectories of bubbles over the waves. However,
their measurements indicated that the kinks appeared somewhat closer to the
water surface, z'= 2-3 cm. The existence of the kinks in the profiles of
air velocity indicate that a jet of high velocity air may sometimes develop
over wavy water in channel flows. To the authors' knowledge, however, with
the possible exception of Sheppard (1952),this phenomenon has not been
Observed with any measurements’ over water in the atmosphere.

Typical measurements of the horizontal distribution of velocity are
shown in Figure 4B. These data are representative of flow in wind tunnels
of rectangular cross-section. It is interesting to note that the boundary
layers associated with the side walls can become rather thick. This thick-
ening had no apparent effect, however, on the development of significant
waves in the channel. The waves still exhibited a nearly linear crest moving
approximately normal to the mean wind direction.

The lines of constant air velocity plotted for a given cross-section
reveal an interesting feature of the channel flow as shown in Figure 5.
Because of a secondary circulation in the tunnel, the lines of constant
velocity are squeezed down in the corners of the cross-section. This has
been observed previously for flow in rectangular ducts (e.g., Schlichting
(1960)). However, the effect appears to become somewhat more pronounced
when fluid flows over a moving boundary in the CSU channel.

The three dimensional structure of the air flow does not visibly affect
the waves generated on the water surface. However, the pressing of the air
moving at higher speeds down along the walls seems to be transmitted to the
horizontal velocity in the water. Measurements of the horizontal distribution
of velocity in moving water show two maxima developing just underneath the
"ears" of the constant velocity curves drawn in Figure 5. Hence, strictly
speaking, the velocity in the air and in water should be written as Uy, Zee
and u(y,z), instead of U(z'), and u(z). However, for the purposes of this
discussion the motion of the air and the water will be treated as two
dimensional.

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F 210.6 meters

y (cm.)

Figure 5. Distribution of air velocity at a given cross-section in
the channel.

296
The vertical profiles for air velocity taken for increasing F along the
center line of the tunnel were found to fit the form:
1/n
a 2 Gly
RO

sR:
WS

where 6 is the thickness of the boundary layer as defined by the value of
z' where U(z') = 0.99U,. Over a wide range of fetch, and for 6.1<U,<13.6
mps the data taken in the channel fit Eq. (1) where n = 4.5. This similarity
distribution is shown in Figure 6. Several typical values of 6 are given
in Table I. The shape corresponding to Eq. (1) with n= 4.5 is frequently
found in wind tunnel data for flow over moderately rough surfaces.

Air Flow Close to the Waves

One of the purposes of this study is to examine the nature of the air
flow close to the water boundary. A key problem in this work centers around
the question of separation of flow to the leeward of the wave crests (Ursell
(1956)). Furthermore, recent theories for energy transfer to the waves from
the air place strong emphasis on the behavior of the region where air
velocity equals the phase speed of waves. For waves generated in the tunnel,
this zone is very close to the water surface, much closer than can be reached
with a fixed probe. That is, the present equipment can only measure average
velocities in the air within about 1 cm of the crests of the highest waves.
To observe the nature of the air motion near U = c, and to trace the
presence of separation, the probe must be placed much closer to the oscillating
water surface than the fixed probe will permit. Therefore, a moving probe
has been designed which will follow the significant waves and maintain ap-
proximately a constant level above the water surface. The schematic picture
of the design, worked out by one of the authors, is shown in Figure 7. The
probe is maintained at a constant level above the water by a servo-driven
mechanism activated by the depth gauges. This system is currently under
construction, and we expect to begin obtaining data from it sometime next
year.

IV. PROPERTIES OF THE WAVES

Over a wide range of air flow which follows the patterns described in
sec. III, only small gravity waves and capillary ripples were generated on
the water standing in the channel. Although the air reached speeds greater
than 12 mps, breaking of waves, in the sense of forming white caps, was not
observed. At high air velocities droplets of spray were observed being shed
from crests of the larger waves, but the waves did not become sharp crested
as seen in "fully developed" seas.

Up to wind speeds of about 3 mps, taken about 20 cm above the water, no
waves appeared on the water surface. However, very small oscillations of ©
the entire water surface could be observed in this range of air flow by
watching variations in reflected light on the water. Above 3 mps, ripples
began to form near the leading edge of the water. These small disturbances

297

Ue (mps)
F (m) 6.06 9.10 13.6

Nm (z- a)/s

d
4 :.
0.6

0.4

0.2

O 0.2 0.4 0.6 0.8 1.0 2

A similarity distribution for the air velocities along the
center section of the tunnel.

298

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Figure 7. The design of the system for positioning instrument probes

near the water surface.

299

had wave lengths of 1 - 3 cm. Their direction of propagation was primarily
normal to the wind direction. As the wind speed increased, the ripples
initially present became larger in amplitude and height. ‘Under the action
of the steady air motion, the waves traveled downstream at an increasing
speed, growing in amplitude and length. For wind speeds in the range

Uoo = 3-6 mps, significant waves were observed to run with crests approximately
normal to the wind direction, with smooth windward surfaces, and rippled lee-
ward surfaces. Above 6 mps, capillary ripples were noted on both the wind-
ward and the leeward sides of the significant waves. At any given point

downstream from the inlet, groups of 5-20 small gravity waves of nearly the
same period passed by. These groups were separated by relatively calm regions
of small ripples having varied periods. The existence of groups of waves
separated by relatively calm water is probably related to interference between
different components of the wave train, giving an appearance of "beats."

The growth with fetch of waves in the channel is reflected in two charac-
teristic lengths, the standard deviation oa and the wave length
The increase with F and Uoo of o and % is show in Figure 8. The effect
of depth is also show in the drawing. Decrease in depth tends to reduce the
wave length, and the standard deviation of the (larger) waves generated at
higher wind speeds. Our data for oa and A were compared to those reported

by Sibul (1955). For a given value of Ugo and d, the results of both these
studies appeared to be essentially the same.

Two characteristic velocities are associated with the water motion. These

are the surface velocity uo, and the phase speed of significant waves c,. The
change with F, Ugo and d of these properties-is shown in Figure 9. Fora
given wind speed, the surface drift remains nearly constant over the range of
d shown, except near the ends of the channel. The wave speed Ce is approxi-
mately independent of depth down to 5.1 cm, but it increases with both Uoo

and F.

Keulegan (1951) found that the ratio of the drift velocity u, to the
wind speed could be correlated with the Reynolds number Reg = u,d/v y,
where Vw is the viscosity of the water. In his calculations, Keulegan
used an air speed averaged over the cross section of his channel, Uavg: Goodwin

(1965) found that U,,, = 0-85 Ugo for the data in the CSU channel. Using this
relation, the values of up have been plotted with Req as shown in Figure 10.
The drift velocities found in this study are correlated satisfactorily in terms
of Reg- Our data fall about 30 percent lower than Keulegan's curve for wavy
water. The difference between these ‘two studies may be accounted for in three
ways. First, if it is assumed that the mass flow of water and air are related
to each other and not the velocities, the momentum ratio, Pwlo/P aUavg

should be used in this correlation. Keulegan's data were taken at sea level
while the CSU measurements were made at nearly 1800 m altitude. If our data
are corrected for the decrease in air density with altitude, they will fall
about 12 percent below Keulegan's curve. The remainder of the difference
between these experiments may be related to (a) the effect of non-uniformities
in up in the y-direction resulting from the nature of the air flow shown in
Figure 5, and (b) the fact that Keulegan used a value of u, averaged over the
length of his channel while our values of uo are taken locally. The effect

of (b) should be small, however, since uo varies little with fetch.

300

@eee00d000

o (cm)

> (cm)

Fetch (meters)
Figure 8. Variation in characteristic lengths of surface waves with fetch
and water depth.

301

0.2

0.15

0.1

u,(mps)

0.05

O 2 4 6 8 0) la 14
Fetch (meters)

0.2

@8eped000

O 2 4 6 8

Figure 9. Variation in characteristic water velocities with fetch and
depth.

302

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303
The Phase Speed and the Drift Current

The phase speed of significant waves Ge as measured with respect to a
fixed point was compared with values calculated from the theory for small
amplitude gravity waves. The theoretical phase speed is:

ust ae
Orne [ ek tanh (a) | (2)

where k = 27/X . In all cases, the values oe 6. were larger than values cal-
culated by Eq. (2). This effect has also been observed by Francis (1951) and
Cox (1958). Francis qualitatively accounted for the deviation by considering
an increase in wave velocity associated with the surface drift, and the fact
that the waves are finite in amplitude. Cox, on the other hand, attributed
the difference to the combined effects of finite amplitude, orbital velocity
of low-frequency wavelets, drift currents, and dynamic effects of the wind.
Cox only analyzed in detail the finite amplitude effect as calculated by
Sekerzh-Zenkovich (1956). Cox found that the finite amplitude effect could
only explain his observed increase in phase velocity for waves larger than

A} = 7cm. The observed differences in phase speed for wavelets of length
smaller than ~% 7 cm could not be accounted for by the influence of finite
amplitudes alone.

For strict comparison to ce Co should be corrected for the mean motion
of the water and not the surface drift since c, should be measured relative
to an average transport in the water. Because the orbital movement of water
particles associated with the waves extends downward to some depth, the sur-
face drift uo is not the proper correction factor. The correction should be
proportional to a weighted average water velocity over some depth below the
surface.

Lilly (1964) has proposed a drift correction for waves traveling on
water at finite depth. Assuming that the vertical profile of the drift cur-
rent is parabolic (laminar flow), and that the waves have infinitesimal ampli-
tudes, Lilly found that

u —
ex mre nee 2 nee = owe #2 Gosh( 2ka) (3)
2 ° 2(ka)2 (kd)sinh(2kd)

For deep water, kd + and Equation (2) implies that the waves travel with
the surface flow only (Gicden Gp 2] Ghee Uy Ye However, for shallow
water, kd +0, and Eq. (3) predicts, as expected, that Ce

The values of cT as calculated by Eq. (3) were compared to the corres-
ponding experimental data, and the results are shown in Figure 11. Experiment
and theory agree within + 15 percent. This error is approximately that
expected on the basis of "experimental errors in estimation of C,, and Cc,

using 2X and Uo. a

304

ic (mps)

Figure ll.

C_(mps)

Comparison between experimental values of the phase speed
of significant waves and values calculated by Lilly's
theory.

Te a, OS eee 3S

305

Systematic deviations between Ce and Cr might be expected since both
the effect of surface tension and of finite wave amplitude were not considered
in deriving Eq. (3). The correction in CT for surface tension in deep water
waves was found to be negligible for the experimental observations. However,
the Stokes correction of Cp for gravity waves of finite amplitude (e.g.,

Lamb (1932)) could vary from 1 percent to 11 percent (increase) if it is
assumed that the amplitude of significant waves is 3.00 (e.g., Sibul
1955)). Thus, Lilly's equation would give values of C, somewhat larger on
the average than the experimental data.

The effect of finite amplitudes in the wind waves may be offset partially
by the influence of turbulence in the water. Dye traces of the motion in the
water indicate that the water flow was turbulent and not laminar. The use
of a turbulent velocity profile having a steeper gradient near the surface
than the parabolic curve but having the same drift velocity at the surface
would result in a smaller correction factor for drift than predicted by

Eigen (3)!

These results indicate that the significant wind waves on the water in
the channel travel relative to a mean drift approximately as gravity waves
of small amplitude. . The waves tend to propagate in this way in spite of the
steady pushing of the moving air.

Shearing Stress on the Water

An important parameter for measuring the action of the wind on the water
is the shearing stress on the water surface, T_. This is often calculated
in terms of the drag coefficient =

B 2
Cre Gal) pu ; (4)

where Ps is the density of the air. An average value of the shearing stress
at the surface tT, can be estimated for the data of the present study by
taking a force balance on the body of air, or the body of water in the chan-
nel at a given time. The shearing stress on the (smooth) walls, on the top
of the tunnel and on the bottom can be estimated in well known ways (@ol8or
Schlicting (1960)). Then the stress t, can be calculated from differences
between the pressure gradient (on the set up of water), and the shear forces
on the walls, bottom and the top. This has been done for our data by Goodwin
(1965). His results are expressed as the average drag coefficient based on
the average air velocity in the tunnel as defined by:

@ os 2
Cy = taf OU ccc (5)
Estimates of C_ as they vary with Usye are shown for two different depths

in Figure 12. For comparison, the data of other investigators, Francis
(1951), Keulegan (1951), and Fitzgerald (1962) are also indicated. Goodwin's
calculations by the force balance technique as applied to the water and the

*(L) ‘ba worg pezeTnoTeo se squepToTsje0o TeooT peseizane
04 Jajer squtod sze[nduet714 sy, °e0ejains 194emM oY4 UO SqUSTOTJJoeoo Seip Uso -eT ainety

. e :

SISNVYS ,
2 /

307

air are not the same, but some differences are expected, which will result
from estimations of the stresses on the walls of the tunnel. The data for
d = 5.1 cm appear to be in reasonably good agreement with the results of
Keulegan and Fitzgerald. However, Francis' results show a steeper increase
in C, with Vave:

It is worth noting here that all of the values of Gs correspond to the
case where the channel bottom is smooth. Fracis' calculation includes the
effect of the average bottom shear _T, as well_as the surface shear in his
value of the drag coefficient (i.e. Cs4p = ac + ™))/ avg" ). This
is equivalent to assuming that the ratio T/ Ha is zero for turbulent
flow in the water. Fitzgerald based his data on the same assumption. How-
ever, Keulegan took t) /t. = 0.25 in estimating C, only. In view of this,
it is not clear why Fitzgerald's data for Cg+p show close agreement with the
results of Keulegan and Goodwin for Cs while Francis' results display a con-
siderably larger variation of Garp with Vave: It is also difficult to under-
stand why Goodwin found such a large change in Gs with depth of the water. At
present, this conclusion cannot’ be checked with the findings of other investi-
gators because their results apply to one water depth only.

Local values of the drag coefficient (Cs' = Ts /P aU." ) can also be
estimated from our data. Although evaluation using the assumption of the
logarithmic profile cannot be applied, the momentum integral technique can
be used (e.g., Schlichting (1960)). The relation for Cg' by this method is:

(- 2 ; i
Gwe ate Srueat i B18 AGBANl (6)
s 1 idx re) |

dx .

—

where 6* is the displacement thickness:
z""

BS + (UU) azn (7)

) is the momentum thickness:

AL
a U(U, - U)dz' , (8)
u 2

d
and z" is the value of (z-d) where U = U,,
To accurately obtain values of C,' from Eq. (6), the slope of U

and values of 6* must be well established. The contribution of Bose)
to the integrals in Eqs. (7) and (8) depends strongly on the region of the

308

vertical profile where the curvature is greatest. In our measurements, this
is poorly defined because the curving portion of U(z') lies too close to the
water for accurate measurement with the fixed probe. Thus, the use of Eq.
(6) with Eqs. (7) and (8) can be expected to give only an order of magnitude
estimate of Cs'

In addition to the problem of using the experimental data in Eqs. (7-9),
the application of a definition for proper air velocity must be considered.
Eqs. (7-9) apply to flow over a solid boundary. When the boundary is moving
and waves are superimposed on this motion, the air speed relative to fixed
coordinates may be an incorrect estimate for U(z'). Two other systems of
velocity coordinates can be used. The air velocity relative to the surface
drift may be a better system, or as Benjamin (1959) has noted, the motion
relative to the phase speed c may be better than the fixed system of
reference. Introduction of either one of these reference velocities will
affect the definitions of 6* , 6, 6, and C,'.

In spite of these difficulties, it is useful as a first approximation
to apply Eqs. (7-9) for evaluating C,'-. Calculations of the local drag co-
efficients based on the data for U(z") were made, and some typical results

for 6* ,6 , and Cs’ are shown in Table I. These results indicate that
C,' decreases somewhat with fetch, but tends to increase with wind speed.

The decrease with fetch is eypileail of the variation in Cs' in the context
of a growing boundary layer over a solid surface.

TABLE I

F(meters) Nominal Average Us, (mes) S(cm) 6&*(cm) 6 (cm) Ge x 10
Air Speed (mps)

2.14 52 4.76 12 200% OC82) S38
4.58 5.20 Wa) | DaEO Yee 2.65
7.03 5.48 T75ok EOE W208: nese
2.14 Tol 7.20 13.07 2.97 Los 5 bG
4.58 7.95 WAG BAGO Bop 92
7.03 8.80 DD SoS BSB 303i
2.14 11.6 10.8 Di Saleh) wean S32
4.58 IQ TOMA fa90) J 235 dae
7.03 12.8 DG 56S oa hoes

Using the_nominal average velocity (0.85 U,, measured at large fetch),
estimates of Ge were calculated by averaging each set of three values of C,'
for the ranges Sof Ugo in Table I. As a comparison to Goodwin's results, the
three estimates of Cs by this method are plotted as triangular points in
Figure 12. The data for Cs by two different calculations check reasonably
well for the 5.1 cm depth of water.

309

The velocity profiles for air were taken primarily at 15.2 cm water
depth, but there should be little change in air flow at 5.1 cm. Ome should
bear in mind, however, that Gs calculated from the data in Table I represents
Ts in a narrow slice along the center section of the channel. Strictly
speaking, Goodwin's estimates of Ts include the variation of Cr in the
y direction. From Figure 5, it is clear that the shear on the water surface
near the maxima in the cross-sectional distribution of air velocity will be
larger than that at the center section. This may account for the differences
between values of Cg at d = 15.2 cm as estimated by Goodwin's method and the
momentum integral technique.

V. WAVE SPECTRA
Autocorrelation Functions and the Frequency Spectra

The time correlations between displacements of the water surface were
calculated from the digitized depth gauge data. The autocorrelation function
R(t) is defined as:

R(t) = é(t,) E(t.) ; T = te" (10)

where €(t,}) and €(ty) are surface displacements taken at the same point
for two different times, t] and t2. The averaging technique in Eq. (10) was
carried out after the method given in Blackman and Tukey (1958).

The function R(1) for waves in the channel was found to exhibit certain
interesting features. A typical example is shown in Figure 13. R(t) was
generally found to oscillate regularly about the R(t) = 0 line with in-
creasing t . Its amplitude decreased sharply initially, but it became
fairly steady at higher values of Tt , though sometimes it varied slowly
as if a lower harmonic was present. The behavior of R (tT) suggests that
there is a tendency for the mutual action of the two fluids to force a
nearly periodic, regular disturbance to develop on the water at a given
fetch in the channel. On the regular waves are superimposed small, random
disturbances which are related to the larger values of R(t) OTS me el ears
This, of course, is precisely the physical picture which developed from
visual observations of the development of the significant waves.

It is well know that the autocorrelation of a periodic function of
period P is another periodic function with period P and a zero mean. Hence,
the period of the significant waves can be estimated from the zero crossings
of the autocorrelation function at large values of fT where the effects of
the random component are small. As seen in Figure 13, the components of
"noise" tend to damp out rapidly, so that the period of the significant
waves also can be calculated approximately from zero crossings of R(t)
over the whole range of t+ . Typical values of 1/P found in this manner
are listed in Table II.

310

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TABLE IL
Case co/ r 1/P a
21 2.74 2.87 2.91
79 GD 4.83 Ah
100 Son 5a F/I 5) 05
113 Dol DD DoS
169 335,24 3631 3,26

The energy spectra were calculated by means of the following relation:

6(f) = R(t) cos2nft dt . (11)
0

The scheme for evaluating the integral in Eq. (11) for a finite record is
given by Blackman and Tukey, (1958). However, instead of the usual technique

of "hanning," it was preferred to obtain a suitable lag window by multiplying
the function R(t) by:

g(t) = (1 + cos = es (12)
m
where Ty, for our data is 3.5. The fading function g(7) has the advantage

of suppressing the periodic component in the autocorrelation function at
large lags without removing any information at short lags. An example of
the faded autocorrelation R'(t) (=Rg) corresponding to the curve of R(t)
is shown in Figure 13.

The spectrum corresponding to the faded autocorrelation R'(t) in
Figure 13 is showm in Figure 14. Note that the tendency toward periodicity
in the wave train also is indicated in this spectrum. Higher harmonics of
the frequency f, for which the energy is maximum appear as indicated in this
Figure. If the waves were perfectly periodic, the idealized spectrum based
on the R(t) curve would develop as spikes of infinite height at n
multiples of f,,. However, because the waves are not truly periodic, and
because of the random components which exist in the signal, the spectrum
actually takes the bumpy shape indicated in Figure 14.

_ _Typical values of f,, are shown with corresponding values of 1/P and
Co/A in Table II. Because of the narrowness of the region containing
most of the energy in the spectra for wind generated waves in the channel,
these three frequencies are approximately equal. Thus, for practical

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313

purposes, the waves in the channel may be characterized essentially by the
properties associated with the significant waves.

The Growth of Waves in the Channel

The frequency spectra calculated at different fetches for the same air
velocity indicate how the waves grow as they move downstream along the
ehannel. A typical set of spectra for increasing fetch is shown in Figure 15.
These curves have been corrected by subtracting out the noise level, and
have been smoothed by a method similar to that discussed by Hidy and Plate
(1965). Near the leading edge of the water (small fetch), the observed
spectrum contains little total energy and is rather sharp. As the waves
travel downstream, the magnitude of the spectral density function increases,
the primary peaks broaden at first, then tend to sharpen up while the values
of fm decrease.

The growth of waves in the range of higher frequencies tends to be
limited as indicated in Figure 15. The complete mechanism for restraining
the growth of the high frequency components is not know. However, it can
be seen that the limitation in growth, in part, can be the result of attain-
ing a balance between gains in energy input from the air and losses by dis-
sipation. The dissipation of energy in small gravity-capillary waves is
probably related to the action of viscosity and sugface tension. The loss
by viscous forces in waves is proportional to (ak)~ (Lamb (1932)) where a
is the amplitude of a wave. As proposed by Longuet-Higgins (1962a), the
‘loss resulting from surface tension can be related to the drain of energy
from larger waves when capillary ripples are formed near the crests of the
larger components. This particular mechanism indicates that the energy loss
is proportional to (a®k3,)- The subscript c refers to the capillary ripple
on the crest of a larger wave. If the interaction between components in the
wave train is a second or higher order effect (e.g., Phillips (1963)), the
action of dissipative processes should balance the input of energy from the
air motion in such a way that the net energy at equilibrium is smaller the
higher the frequency range. This seems to be suggested in the behavior of
the spectra shown in Figure 15.

It is interesting to note that the growth of components in the lower
frequency range, say f < 3.5 cps in Figure 15, is approximately exponential.
Qualitatively, this type of growth has been predicted in the recent shearing
flow theories of Miles (see, for example, Miles (1960)).

Similarity Shape of the Spectrum
An important feature which was also exhibited by many of the spectra for
the channel waves was the tendency for growth in such a way that a similarity
shape in the spectral density function is maintained. The frequency spectrum
can be expressed, with Eq. (11) in normalized form, as:

(12)

314

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(sd9)4
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(90s

315

fm = W(£/Eq) (13)
o2

where denotes a dimensionless quantity representing a "uriversal"
spectral density function.

Typical smoothed spectra which have been smoothed and corrected for
noise level after Hidy and Plate (1965) are plotted in a form corresponding
to Eq. (12) in Figure 16. The spectra in Figure 16 serve to define the
similarity function quite well. The conditions of Ugo, F and d for these
spectra are shown in Table III along with the values of o , f, and a
In general, it was found that the channel data followed wW quite satis-
factorily for the range 6 <U < 15 mps, and for 3 <F< 12 meters.

TABLE III
Case d(meters) U__ (mps) F (meters) fo} 10° f (cps) ®
fore) m m
(meters)
163 0.0254 6.10 552k 0.131 4.83 2.90x10" +
LS ORO2; O57 10.7 0.767 2,36 7.99x10_,
188 0.102 IL 8.16 13)6 98 2.97%10_.,
1.92) 0.0508 10.7 Lilo 0.538 Do 33} 5.53x10_,
208 0.0508 9.15 7.86 0.614 2.48 8.65x10_,
212 0.102 IO 7/ 5.74 0.457 35 17/ 4.15x10

According to Phillips (1958a), on dimensional grounds, the equilibrium
or saturation region in the high frequency region of the spectra for gravity
waves should follow the f-7 rule. In contrast, it has been suggested by
Hicks (see, for exemple, Phillips (1958b)) that the pure capillary spectrum
should follow an f-!/3 rule. As indicated in Figure 16, the dimensionless
spectra for waves in the channel tend to follow the f°’ rule over approxi-
mately two decades in the high frequency range. In the highest frequency
ratios, there is a tendency for some of the spectra to develop a slope less
than -5. Capillary wave behavior should begin to appear above f = 13 eps
in the frequency spectra. Only two cases, numbers 163 and 188 as show in
Figure 16, actually reach this range. For case 163, capillary waves should
appear for 5// 2 = 2.7 to 3.0, while, for case 188, £/e =6.8 to 7.0. Thus,
in Figure 18, these two examples may display the beginnings of a transition
to the #771/3 range.

316

10°! 10
f/f

Figure 16. The similarity spectrum for wind generated waves in the channel.
After Hidy and Plate (1965).

317

Unfortunately, the existence of the 77/3 rule cannot be verified
generally in these results because the highest frequencies which can be
resolved with some accuracy in the computation scheme used, lie around
15 eps. Therefore, these data cannot provide conclusive evidence for the
existence of an equilibrium range in capillary waves.

VI. SOME COMPARISONS BETWEEN WAVES IN THE CHANNEL AND OCEAN WAVES

This study was not intended specifically to model waves at sea. Never-
theless, it is worthwhile to see where the data taken for the channel waves
fit into the overall picture of wind generated waves.

The properties of the wind waves generated in the channel easily can be
placed in perspective with much larger scale conditions by using well know
pictures of wave behavior. As an illustration, the channel data have been
plotted schematically in two "scaling’’ drawings of Hicks (1963), as shown
in Figures 17A and B. In Figure 17A, the logarithm of the standard deviation
is plotted against the logarithm of fetch. The numbers near curves or points
refer to wind speeds taken in the field at somewhat different anemometer
heights. The values of wind speed for our data correspond to Ugo: Our
data for waves in the channel fit nicely into the extreme region of short
fetch of this figure. Similarly, Figure 17B shows the variation in f,, with
fetch. Again our results for small waves correspond to field data taken at
very small fetch.

Another way of illustrating where the wave data for the channel fit into
the geophysical picture comes from the correlation curves of Wiegel (1963).
On dimensional grounds, the mean properties of wind generated waves (on
deep water) should be related to a Froude parameter with the characteristic
length being the fetch. This correlation shows essentially the variation
in standard deviation, length of significant waves, and frequently with
fetch.

Wiegel's results include data taken over an extremely wide range of
conditions, which include results of some laboratory experiments and field
studies on lakes and on the ocean.

A plot of a number of values of o , a and f for the waves in the
CSU channel is show in Figure 18 along with average curves estimated from
Wiegel's Figure 6-5-7. The averaged curve for the correlation of standard
deviation (or 0.33 times the significant wave height Hy /3) falls approxi-
mately along the mean of the data from this study. Furthermore, the lines
of maxima and minima in o , as estimated from Wiegel's correlation, encom-
pass all of our values. It is interesting to note, however, that there is
a systematic deviation with Uo, for the parameter go/U,,* in Figure 18.
This indicates that the Froude criterion gF/Ug9° cannot be the only parameter
for modeling the mean displacement of the water surface. Furthermore, it is
clear from the data in Figure 8 and the points for high wind velocity in
Figure 18 that the water depth should be included in correlating heights for
waves moving at finite depth in channels.

318

-2.0 | 7

b Hip GE This Study
= thei : eae.

S -25 are = Burling

-40

Log F (meters)

Figure 17A. Properties of small wind waves. Standard deviation.
After Hicks (1963).

HB This Study

Kinsman

4-8
S lom/sec

Log, F (meters)

178. Properties of small wind waves. Frequency of significant
waves. After Hicks (1963).

320

Figure 18.

I
10 af
@ .
Froude number correlation of properties of wind waves.
Wiegel (1963).

After

SS SE TT _

321

The data for the dimensionless wave length of significant waves in the
channel fit Wiegel's correlation curve satisfactorily. There is a tendency,
however, for our points to lie somewhat above the estimate of Wiegel's line.

A plot of the wind tunnel data for f, indicates that Wiegel's curve is
again approached. However, in this case, our data lie somewhat below the
estimated correlation curve. This is somewhat misleading, however, since
comparatively little data is presented for this value (in terms of P) in
Wiegel's paper.

The dimensionless energy spectrum derived by Eq. (12) can be compared
to the dimensionless form discussed by Bretschneider (1963a and 1963b),
provided that it is assumed that . e/a is approximately equal to
0.34 4, where $, is the maximum value of the spectral density
function (see, for examples, the values in Table III). This has been done
and the comparison is shown in Figure 19. Bretschneider's estimate for
the Neumann spectrum corresponding to fully developed sea is shown along
with our dimensionless spectrum. For further comparison, typical data of
Burling as calculated by Bretschneider are shown in Figure 19. As might be
expected, the dimensionless spectrum for waves at very short fetch in the
channel is narrower than Burling's data taken at F = 1000 meters and the
Neumann spectrum for nearly infinite fetch. The three dimensional character
of waves developing on the surface of the oceans or lakes, of course, cannot
be reproduced in the channel.

Acknowledgement

The wind-water tunnel at Colorado State University was constructed under
a matching grant from the National Science Foundation. This work was supported
by the National Science Foundation in connection with its grant to Colorado
State University, and its contract with the National Center for Atmospheric
Research. The authors are grateful to R. Biro, C. Goodwin, Hosein Shokouh,
C. Yang and many others for their efforts in carrying out the experimental
program and analyzing the extensive data of this study.

322

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323
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of California, Berkeley, California.

Some engineering aspects of wave
spectra, in Ocean Wave Spectra.
Prentice-Hall, In Inc., -, Englewood
Cliffs, N. J-, 357 pp-

Wave generation by wind, in

Surveys in Mechanics. Cambridge

University Press, 475 pp.

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ON THE INSTABILITY OF EKMAN BOUNDARY FLOW

Douglas K. Lilly
National Center for Atmospheric Research
Boulder, Colorado

Note: The full text of this address has been submitted for
publication in the Journal of the Atmospheric Sciences.

327

329
ABSTRACT

Experiments by Faller (1963) revealed the existence and nature of a shear-
ing instability in laminar Fkman boundary flow. Barcilon (1964) formated
the appropriate linear stability equations and obtained partial and approximate
solutions, based on the large Reynolds number methods appropriate to other
related problems. Numerical solutions of the perturbation differential equa-
tions have been obtained for a large number of variations of the three variable
parameters, «a , the dimensionless wave number, e , the angle of the instability
bands relative to the geostrophic flow, and R, the Reynolds number. Solutions
are presented for the eigenvalues, the real and imaginary velocity components,
and for the eigenvectors, the actual modes of the normal, parallel, and vertical
velocity components.

Two essentially different unstable modes are found, only one of which cor-
responds to Faller's measured results. This solution, which we call the normal
mode, derives its energy from the mean flow component normal to the bands. It
is characterized by a band orientation lying between the geostrophic and surface
flow directions, i.e. € is positive and less than 45°. The critical Reynolds
number is about 110, compared to Faller's measured value of 125, and the critical
values of a,c and the small real wave velocity are in good agreement with
the experiments. This instability is adequately explained (Faller, 1963, also
see Stuart, 1955, for a related problem) by inviscid theory, and is associated
with the principal point of inflection in the normal mean velocity profile.

The other instability mode obtains its energy from the mean flow shear com-
‘ponent parallel to the instability bands, and we therefore denote it as the
parallel mode. At the critical point the bands have an angle e <0 , that
is a direction outward from the geostrophic flow, and a wave length about
double that of the normal mode bands. The critical Reynolds number is about
60, thus in the experimental realizations the parallel’ mode should be present
before and together with the normal mode. Faller apparently observed this
instability qualitatively but erratically. His dye visualization method was
probably most responsive to nearly stationary bands, while the parallel mode
bands have a large inward real wave velocity.

Further analysis of the dynamics of the parallel mode reveal similarities
and differences with other types of hydrodynamic instability. Rotation is of
critical importance, as can be determin’ d by comparison of numerical solutions
or the perturbation equations with and without the Coriolis terms. Analytic
solutions can be obtained from a simplified version of the perturbation equa-
tions, which compare fairly closely with the complete numerical solutions.

A similarity can be seen to exist with baroclinic instability for cases of
short wave length and neutral static stability.

Comparisons between the instabilities present in laminar boundary flow
with the finite amplitude disturbances existing in a real turbulent boundary
layer are at best qualitative. It may be postulated, however, that either or
both of the modes of energy transfer described here could form the principal
energy-containing eddies of the atmospheric planetary boundary layer.

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Tihs ee

FEDERAL RESEARCH PROGRAMS IN AIR
SEA INTERACTION

Feodor Ostapoff
U. S. Weather Bureau
Washington, D. C.

331

BLAS SMA

RORARAEE, AAAS:

333

The conclusion of a conference of this sort may be an appropriate time
to discuss the role of the Federal Government in air-sea interaction research,
since, as you know, considerable government attention is now directed to this
topic.

Let me say first, that, not long ago, air-sea interaction research had
not been identified as such, but was carried out under either the National
Atmospheric Science Program or the National Oceanography Program. As a matter
of fact, this was the first year that an item named "air-sea interaction
research" has been carried in the Atmospheric Science Program.

Why this recent burst of activities and special emphasis? For a long
time man concerned himself with problems ofthe sea. Obviously, phenomena at
the sea surface were always of greatest interest, and these are problems which,
in one way or another, are the direct result of meteorological events. For
example, water waves aroused the interest of the most brilliant minds of the
last 200 years. Names like Lagrange, Cauchy, Poisson, Airy, Stokes, Kelvin,
Euler, and others come to mind. Today, waves are still subjects of intense
investigations. Other examples could be cited such as concepts of the oceanic
circulation. As early as 1686, reference was made by Halley that the surface
circulition of the ocean iswind-driven. It is most difficult to conceive for
the ocean such a simple model of meridional circulation such as the Hadley
cell in the atmosphere. For the ocean is heated and cooled at the surface,
or, in other words its heat sources and sinks are essentially at the same
geopotential. Meridional convection processes seem to be rather inefficient.
On the other hand, when about 80 years ago the first deep-sea temperatures
were taken in the tropical ocean by the Challenger Expedition and very low
temperatures were observed at great depth, it was immediately reasoned that
this water is of polar origin and must have descended to a great depth due to
cooling at the surface. The thermohaline circulation, which for a number of
years was subject to little attention, is now being restudied, for example

by Dr.K . Bryan.

This merely illustrates the difficulties we face in defining air-sea
interaction research. It is quite possible -- and this has been done -- to
include in such a category almost all of meteorology and oceanography. How-
ever, for the purpose of this discussion we like to narrow the definition of
air-sea interaction because we, in the government, find it most advantageous
for administrative purposes to consider in a program only the direct aspects
of air-sea interactions. Of course, conferences such as this should not
adopt this narrow parochial view and should discuss broad problem areas.

Historically, the Second World War had a profound mpact on both
oceanography and meteorology. Meteorology then advanced with the help of
modern high-speed computers to such an extent that realistic models of the
general circulation could be developed and tested and the entire prediction
service operated on a more Objective basis. The degree of sophistication
incorporated in these models increased to a level that soon effects of the

334

underlying surface become important. Increased numbers of observations and
entire new systems were introduced. And again, high-speed computers made
it possible to digest the wealth of information. Thus, theory and observa-
tion began to exist on an equal footing. Indeed, it is now realized that

a new approach to the observational program is in order.

One of the problems facing us, an urgent one in meteorology, is how to
obtain new and more accurate information on energy exchanges between the
ocean and the atmosphere and vice versa. This problem was investigated in
the early 1960's by a joint panel on air-sea interaction of the National
Academy of Sciences Committee on Oceanography and the National Academy of
Sciences Committee on Atmospheric Sciences. The Chairman of this Panel is
Dr. Benton. The work of the Panel culminated in National Academy of Science's
Publication No. 983, 1962. This report outlined eloquently the problem areas
of air-sea interaction and formulated six recommendations which, if implemented,
will lead (in the Academy's opinion) to a new approach to the general problem.
Today I would like to stress only one of the six recommendations which
introduces the concept of "area studies."

An area study is in effect what meteorologists would call a synoptic
scale investigation over a suitable and carefully selected oceanic area where
a number of observational techniques can be employed and intercompared. The
problem of energy exchanges is considered of such complex nature, and the
available methods still not completely adequate (perhaps with the exception
of direct Reynolds stress measurements and vertical heat flux determinations)
that it seems necessary to plan and to conduct a thoroughly redundant experi-
ment in order to gain some confidence in any oneof the existing techniques.
The study area will be bounded by a line of meteorological observations at
its circumference in order to evaluate line integrals and apply continuity
considerations as independent checks on other indirect techniques.

The Federal Government responded to the Academy report when the Inter-
agency Committee on Oceanography and the Interdepartmental Committee on
Atmospheric Sciences jointly appointed an Ad Hoc Panel on air-sea interaction
chaired by Dr. Jacobs, Director of the National Oceanographic Data Center.

In 1963 the Panel made several recommendations one of which has been acted
upon by the Interagency Committee on Oceanography and the Interdepartmental
Committee on Atmospheric Sciences. Subsequently, the Federal Council for
Science and Technology agreed to assign special responsibility for the air-
sea interaction research to the Department of Commerce. The following letter
records the assigning of this responsibility:

6 March 1964
MEMORANDUM FOR:

J. Herbert Hollomon, Chairman, ICAS
James H. Wakelin, Chairman, ICO
SUBJECT: Air-Sea Interaction Program

At its meeting on February 25, 1964, the Federal Council discussed
questions related to the responsibility of the Department of Commerce for

335

‘the air-sea interaction program. The Council supported the designation
of special responsibility in this area to DOC as an experiment to be
reviewed at the end of one year. The Council further agreed to the
following specific statement of this responsibility as it appears in
the minutes;

"Planning Responsibilities for Interagency Air/Sea Interaction
Research

1. The Department of Commerce is responsible for coordinated
planning of the Federally-supported air/sea interaction program.

2. This program will be developed in consultation with the
participating agencies through their representatives on a joint
ICO/ICAS air/sea interaction panel.

3. The program developed by DOC will be submitted jointly to
ICO/ICAS for review.

4. These air/sea interaction plans will be submitted to the
Federal Council for Science and Technology for endorsement as components
of National Programs in Oceanography and Atmospheric Sciences.

5- Agencies are responsible for funding, conducting, and other-
wise managing their portions of the program, consistent with their
statutory missions. Such portions are those either initially proposed
by an agency for support, or added by mutual agreement between the
agency, DOC, and the joint panel.

6. DOC will undertake funding needed to fill out the remaining
important components of the total program plan consistent with its
statutory missions. Area studies of transfer processes represent one
such compoment."

The intention of the Council is to stimulate activity and provide
effective leadership and coordination of the air/sea interaction program.
I am glad to endorse the Council's statement, which should serve to
amplify Dr. Wiesner's earlier memoranda on the subject and should clarify
the relationship of DOC to the other agencies in this area.

/s/ Don
Donald F. Hornig
The Department of Commerce responded organizationally to the assigned
responsibility by establishing a Sea-Air Interaction Laboratory (SAIL) which

is operated jointly by the Weather Bureau and Coast & Geodetic Survey.
Furthermore, in order to provide a mechanism for coordination of the program

336

and for consultation in the development of a program a permanent Joint rco/
ICAS Air-Sea Interaction Panel was established in May 1964.

At that time Dr. Frank Gifford was the Chairman of this Panel. When
Dr. J. Spar became Director of Meteorological Research of the Weather Bureau
he then assumed the Chairmanship. The members of the Panel are:

Thomas S. Austin: - Dept. of Interior, Bureau of Commercial Fisheries

Dr. Woodrow C. Jacobs - National Oceanographic Data Center

Armold B. Joseph - Atomic Energy Commission

Dr. Donald Martineau - Dept. of Defense

Ledr. R. M. Morse = Dept. of the Treasury, Coast Guard

Dr. R. V. Thomann - Delaware Estuary Comprehensive Study - Dept. of
Health, Education and Welfare

Dr. Fred White - National Science Foundation

In order to proceed with the development of a federal air-sea interaction
program, it was first necessary to go through the atmospheric sciences program
and the oceanography program in order to identify air-sea interaction research
or contributing activities in these programs. Therefore, we attempted to
define what we mean by air-sea interaction research. The following definition
has been adopted. Research projects must satisfy one or more aspects of the
following goals of principal objectives: (1) The development of a sound body
of theoretical and empirical knowledge relative to the processes by which
energy transfer and transformation occur in the ocean - atmosphere boundary
layers across the entire spectrum of time and space scales and, (2) The
identification of those processes which are of critical importance in the
prediction of the system's future states, with emphasis on understanding
those processes which are susceptible to control by artificial means.

There are several ways in which one could further categorize research
activities and bring some order into the field. One could discuss air-sea
problems in terms of micro-scale, meso-scale, and large-scale problems, which
may be arbitrary but convenient. However, these terms do not mean the same
thing to everybody. We have prepared an Interim Report on federally-supported
air-sea interaction research in FY-64, FY-65, FY-66. This report summarizes
the present program. In the report we speak in terms of (1) dynamics of the
boundary layer and turbulence, (2) physics and chemistry of air-sea interface,
and (3) large-scale field studies. The last category really encompasses
elements of the first two. However, special efforts are required to mount
such projects. They may require coordination of a special kind because of
use of different systems, platforms, and techniques. They may also require
interagency coordination and close cooperation with universities and private
research institutions.

Category I, the "boundary layer dynamics and turbulence" contains
subject matters such as:

- _methods to measure vertical eddy fluxes in the turbulent boundary
layer

- theories to explain the mechanisms of turbulent flux and its
relation to the structure of theboundary layer

- Jaboratory experiments and mathematical models to study boundary
layer processes

» empirical formulae to relate exchange of momentum, sensible heat,

and water vapor to the structure of the atmosphere, hydrosphere,
and the interface, and

- circulation models of the atmosphere and hydrosphere which include
interaction mechanisms

Perhaps there is now a little time available to indicate briefly what
is being done in the government in this field, leaving out significant and
probably most important contributions from universities. Some results of
these studies have been reported in this conference. The main federal
support for university work comes from the National Science Foundation and
the Office of Naval Research.

Perhaps the Navy's sea-air interaction program is the largest of its
kind in the government. One large study considers the generation of ocean
waves and its relationship tothe mechanisms of energy transfers from air to
sea. Mainly, Reynold's stresses are measured directly under a great
variety of conditions. This investigation includes also heat flux measure-
ments. There is a significant study under way to investigate the mechanisms
of coupling the air =- ice - water system and its relationship to the heat
budget of the Arctic Ocean. The objective of the program is to predict or
improve the prediction of ice drift, of pressure - ridge formation, of the
stability of leads in the ice as well as the formation, growth, deformation
and disintegration of sea ice. Computer programs are designed to test
statistically and dynamically three-dimensional models of ocean circulation,
the application of environmental parameters representing fluxes of the air-
sea interface as they relate to macro-scale oceanic processes, and the
influence of persistence, climatic, and transient synoptic processes; and
frictional and heat energy exchange on the behavior of the ocean. Further-
more, a project deals with the energy budget of an air-water column extending
from about 16 meters below to 45 meters above the interface. Micro-scale
fluxes of momentum, heat, and humidity are measured at a fixed tower. The
results will provide insight into the mechanisms of physical processes on the
micro-scale.

The Public Health Service has initiated and is conducting a comprehensive
program on water pollution control. This program is carried out over the
Great Lakes and encompasses a system of buoys, floats, and surface observations.

338

The Dept. of the Interior, in particular the Bureau of Commercial
Fisheries, is very mich interested in air-sea interactions as they relate
to heating, cooling, upwelling, advection, convergences and the distribution
of fish, in particular the tuna. For this purpose the Bureau is assembling
month-by-month historical mean sea-surface charts, is correlating and
analyzing sea-level atmospheric pressures, sea-water temperatures, and winds
at the 20 stations in the North Pacific and studies the relationship between
winds, currents and other vector quantities. Studies are being made of heat
exchange at the air-sea interface to develop prediction techniques for eastern
pacific temperate tuna.

There are many more projects going on in the Federal Government which
deal with this category. However, due to the limitations of time we are not
now able to mention them all. It is important to remember that these projects
are conducted within the agencies. A large number of projects are contracted
with universities and private research institutions.

The following agencies are conducting or supporting research in air-sea
interaction:

Department of Defense:

Department of the Army
Department of the Navy

Department of Health, Education and Welfare:
Public Health Service
Department of Interior:

Bureau Commercial Fisheries;
Bureau of Sports Fisheries

Department of the Treasury:
Coast Guard

Atomic Energy Commission

National Science Foundation

Department of Commerce:

Coast & Geodetic Survey
Weather Bureau

339

Category II, namely, “physics and chemistry at the air-sea interface"
is mostly being done outside the government under contract support. Therefore,

I will not go into any details of these projects.

Category III which we defined as "large-scale field studies" contains a
number of activities. As a matter of fact we could place into this category
any major investigation of any geographical area. However, I would like to
mention at this point that to our knowledge no study is under way at this
time which would satisfy the criterion for an "area study" as defined previously.
Of course, in the category of large-scale field studies we would describe the
work done during the Indian Ocean Expedition, in particular, its marine
meteorological and surface oceanographic activities.

The Navy has a large program in the Arctic which could be considered as
a large-scale study. Observations are being obtained from aircraft, satellite,
and submarine reconnaisance. The data obtained will reveal time and space
distributions of quantities of ice and open water, discrete open water
features, and distribution of pressure ridges and ice thickness throughout
the Arctic Basin. The data will be used in the development of mathematical
models which relate broad scale thermal regimes and motion fields of both
media to the ice features.

The Department of Interior has initiated a trade-wind zone program which
represents a quantitative study of the seasonal processes as they relate to
the Pacific tradewind system, particularly at the air-sea interface. The
area of study here is 10 degrees to 30 degrees North and 140 degrees to 180
degrees West. This effort will also attempt to define seasonal displacement
of the system and lateral changes of heat fluxes across the air-sea interface
and to determine climatic regimes and location and strength of the trade-wind
systems.

The Department of Commerce (in addition to Dr. Bryan's work in the
theory of ocean circulation, which includes both a wind-driven and thermohaline
mode) is mounting now a program along the East Coast which may have some elements
of air-sea interaction. Basically an observational program, hydrographic time
eross section measurements are planned off South Carolina. The analysis will
involve possible meteorological influences on oceanographic parameters. The
Weather Bureau increasingly will also make use of its Research Flight Facility,
utilizing instrumented aircraft in sampling the atmospheric boundary layer
over the ocean in all kinds of meteorological conditions. Further, the Weather
Bureau will also continue its storm surge research.

Where do we go from here? It is undeniable that considerable efforts
are being made in air-sea interaction research, and at considerable expendi-
ture. Indeed, the Federal Government presently supports this type of research
at about a level of $6,000,000 a year. This amount is spent for in-house
research activities as well as for contracts and grants to universities and
research institutions.

340

It seems now that we need a comprehensive and coordinated program. We
are attempting now to develop a program on scientific technical grounds. The
next step is to evaluate the existing program as presently composed of many
projects in the various departments and agencies. Finally, the existing
program needs to be evaluated in terms of what a program ought to be: missing
or weak activities need emphasis and strengthening, and perhaps too much
emphasis has been placed in other areas and may represent undesirable duplica-
tion. It is also conceivable that we may wantto promote "duplication" in
some areas in order to provide necessary competition and supplementation.

The Sea-Air Interaction Laboratory is now engaged in the development
of the federal program plan. Hopefully, this will be accomplished by the
end of this year. In my opinion, this conference contributed significantly
in clarifying the research problem areas we face in this important field.

341
APPENDICES
I. PROGRAM

SEA-ATR INTERACTION CONFERENCE
TALLAHASSEE, FLORIDA
FEBRUARY 23-25, 1965

The conference was sponsored by the Department of Meteorology and the
Oceanographic Institute, Florida State University and by the Sea-Air
Interaction Laboratory of the Department of Commerce (a joint laboratory of
the U. S. Weather Bureau and the U. S. Coast and Geodetic Survey).

SESSION I

3:00 PM, Tuesday, February 23
Room 101, Mathematics-Meteorology Building
Florida State University

Welcome: Dr. Karl Dittmer
Vice President for Academic Affairs
Florida State University

ROLE OF SEA-AIR INTERACTION IN THE ATMOSPHERIC CIRCULATION

Chairman: Dr. N. E. La Seur, Professor of Meteorology, Department of
Meteorology, Florida State University.

"Air-Sea Exchange as a Factor in Synoptic-Scale Meteorology in Middle
Latitudes." Dr. Jerome Spar, Director, Meteorological Research, U. S.
Weather Bureau, Washington, D. C.

"The Three-Dimensional Ocean Circulation Driven by Density Gradients in
an Enclosed Basin," by Dr. Kirk Bryan, Geophysical Fluid Dynamics Laboratory,
U. S. Weather Bureau, Washington, D. C.

SESSION IT

8:30 PM, Tuesday, February 23
Wakulla Springs Hotel

Chairman: Dr. Fred D. White, Program Director for Meteorology, National
Science Foundation, Washington, D. C.

"On the Present State of Knowledge in Air-Sea Boundary Layer Problems."
Dr. H. U. Roll, NSF Senior Foreign Scientist and Visiting Professor of
Meteorology, Florida State University. (Dr. Roll is now President of the
German Hydrographic Institute. )

342
SESSION III

9:00 AM, Wednesday, February 24
Wakulla Springs Hotel

THE ROLE OF SEA-ATR INTERACTION IN TROPICAL METEOROLOGY

Chairman: Dr. Herbert Riehl, Head, Department of Atmospheric Sciences,
Colorado State University.

"A Survey of the Role of Sea-Air Interaction in Tropical Meteorology."
Dr. Joanne Malkus Simpson, Director of Project STORM FURY, U. S. Weather
Bureau, Washington, D.C.

"Sensible and Latent Heat Exchange in Low Latitude Synoptic-Scale Systems."
Dr. Michael Garstang, Associate Professor of Meteorology, Oceanographic
Institute and Department of Meteorology, Florida State University.

"Intensity of Hurricanes in Relation to Sea Surface Energy Exchange." Irving
Perlroth, National Oceanographic Data Center, Washington, D.C.

"The Gulf of Mexico after HILDA (Preliminary Results)." Dr. Dale F. Leipper,
Professor of Oceanography, Department of Oceanography and Meteorology, Texas
A&M University.

"Evidence of Surface Cooling Due to Typhoons," by Dr. C. L. Jordan, Professor
of Meteorology, Department of Meteorology, Florida State University.

"The Modification of Water Temperatures by Hurricane CARLA," by Drs. Robert
E. Stevenson, Associate Professor of Meteorology, and Reed S. Armstrong,
Graduate Assistant, Oceanographic Institute, Florida State University.

SESSION IV

2:30 PM, Wednesday, February 24
Wakulla Springs Hotel

THE ROLE OF SEA-AIR INTERACTION IN TROPICAL METEOROLOGY (Continued) AND
LARGE-SCALE ASPECTS OF AIR-SEA INTERCHANGE

Chairman: Dr. Woodrow C. Jacobs, Director, National Oceanographic Data
Center, Washington, D.C. :

"On the Low Level Thermal Stratification of the Monsoon Air over the Arabian
Sea and its Connection to the Water Temperature Field." Dr. José A. Colén,
Meteorologist in Charge, U. S. Weather Bureau, San Juan, P. R.

"A Low-Level Jet produced by Air, Sea, md Land Interactions." Dr. Andrew F.
Bunker, Woods Hole Oceanographic Institution.

343

"y. S. Fleet Numerical Weather Facility Activities Relating to Sea-Air
Interactions on a Synoptic Scale." Comdr. W. E. Hubert, U. S. Fleet Numerical
Weather Facility, U. S. Naval Postgraduate School, Monterey, California.

"Synoptic Scale Heat Exchange and its Relations to Weather," Dr. Taivo
Laevastu, U. S. Fleet Numerical Weather Facility, U. S. Naval Postgraduate
School, Monterey, California.

SESSION V

9:00 AM, Thursday, February 25
Room 201, Education Building
Florida State University

RESEARCH IN AIR-SEA INTERACTION PROBLEMS

Chairman: Dr. George S. Benton, Chairman, Department of Mechanics, Johns
Hopkins University.

"Laboratory Studies of Wind Action on Water Standing in a Channel."
Dr. George M. Hidy, National Center for Atmospheric Research, Boulder,
Colorado.

"On the Instability of Ekman Boundary Flow." Dr. Douglas K. Lilly, National
Center for Atmospheric Research, Boulder, Colorado.

"Federal Research Programs in Air-Sea Interaction." Feodor Ostapoff, Acting
Director, Sea-Air Interaction Laboratory, U. S. Department of Commerce,
Washington, D. C.

3h
APPENDICES

II. PARTICIPANTS

Thomas S. Austin, Bureau of Commercial Fisheries, Washington, D.C.

Tim P. Barnett, Navy Oceanographic Office, Washington, D. C.

George S. Benton, Johns Hopkins University, Baltimore, Md.

Sherman Betts, Interagency Committee on Atmospheric Science, Washington, D. C.

Walter A. Bohan, The Walter A. Bohan Company, Park Ridge, Ill.

Andrew F. Bunker, Woods Hole Oceanographic Institution, Woods Hole, Mass.

Kirk Bryan, U.S. Weather Bureau, Washington, D. C.

José A. Colon, U. S. Weather Bureau, San Juan, Puerto Rico

Mariano A. Estoque, University of Miami, Coral Gables, Fla. (on leave from
the University of Hawaii)

George M. Hidy, National Center for Atmospheric Research, Boulder, Colo.

William E. Hubert, Comdr., USN, U. S. Navy Numerical Weather Facility,
Monterey, Calif.

Woodrow C. Jacobs, National Oceanographic Data Center, Washington, D. C.

Arnold B. Joseph, Atomic Energy Commission, Washington, D. C.

A. D. Kirwan, Jr., New York University, New York, N. Y.

Taivo Laevastu, U. S. Navy Numerical Weather Facility, Monterey, Calif.
Dale F. Leipper, Department of Oceanography and Meteorology, Texas A and M
University, College Station, Texas
Douglas K. Lilly, National Center for Atmospheric Research, Boulder, Colo.

Donald P. Martineau, Office of Naval Research, Washington, D. C.

Banner I. Miller, U. S. Weather Bureau, Miami, Fla.

Richard Morse, Comdr. U.S.C.G., U. S. Coast Guard Oceanographic Unit,
Washington, D. C.

Thomas H. R. O'Neill, Comdr. U.S.N., Naval Weather Service, Washington, D. C.

A. H. Oshiver, U. S. Coast and Geodetic Survey, Washington, D. C.

Feodor Ostapoff, Sea Air Interaction Laboratory, U. S. Department of Commerce,
Washington, D. C.

Irving Perlroth, National Oceanographic Data Center, Washington, D. C.
Robert 0. Reid, Department of Oceanography and Meteorology, Texas A and M
University, College Station, Texas

Joanne Simpson, U. S. Weather Bureau, Washington, D. C.

S. Fred Singer, University of Miami, Coral Gables, Fla.

R. Bruce Snyder, U. S. Navy Weather Research Facility, Norfolk, Va.
Jerome Spar, U. S. Weather Bureau, Washington, D. C.

Dee F. Taylor, U. S. Forest Service, Macon, Ga.

Fred D. White, National Science Foundation, Washington, D. C.

John Young, ESSO Research and Engineering,.-Linden, N. J.

Bernard D. Zetler, U. S. Coast and Geodetic Survey, Washington, D. C.

345

Florida State University Participants

Reed S. Armstrong, Graduate Assistant, Oceanographic Institute

Philip A. Calabrese, U. S. Weather Bureau Scholarship Student, Department
of Meteorology

Albert W. Collier, Director, Oceanographic Institute

Gordon A. Dean, Research Associate, Department of Meteorology

George E. Fisher, U. S. Weather Bureau Scholarship Student, Department
of Meteorology

Michael Garstang, Research Associate in the Oceanographic Institute and
Assistant Professor of Meteorology

Joseph H. Golden, Graduate Assistant, Department of Meteorology

John C. Gille, Assistant Professor of Meteorology

Seymour L. Hess, Professor of Meteorology

Charles L. Jordan, Professor of Meteorology

Noel E. LaSeur, Professor of Meteorology

Hans U. Roll, N.S.F. Senior Foreign Scientist, Visiting Professor of
Meteorology

Robert E. Stevenson, Research Associate in the Oceanographic Institute
and Associate Professor of Meteorology

Detlef A. Warnke, Research Associate, Oceanographic Institute

Victor Wiggert, Graduate Assistant, Department of Meteorology

William L. Woodley, Graduate Assistant, Department of Meteorology

Edward J. Zipser, Research Associate, Department of Meteorology

USCOMM-WB-DC

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