Section I:Selected Topics in Marine

and Coastal Climatology

by James L. Wise and Lynn D. Leslie

I-1

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Figure 1. MMS Lease Sale Areas

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Aleutian””Islands

Figure 2. Place Names Map

l-3

Bathymetry in Meters

0 1 0 0 2 0 0 3 0 0

MILES

u”KILOMETERS

Source: USGS 1:2,500,000 Open FileBathymetry Series, 76-821, 76-822,76-823, Glen Schumaker.

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Figure 3. Bathymetry

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Currents of the Bering Sea

North Aleutian Shelf

The primary flow of water into the BeringSea originates at Unimak Pass. The sourcewaterofthisflow istheAlaskanCoastalCurrent,from south of the Aleutians. Within the pass andnorth of Unimak Island, much of the coastal cur-rent is entrained into the wind-driven flow alongthe north Aleutian coast. Typically, this currentflows to the northeast into Bristol Bay in thedirection of the prevailing wind, followingbathymetry contours along the coast. At times,the north Aleutian coastal current will undergo areversal in direction due to changes in the large-scale and mesoscale wind direction. Becausewinds are highly variable, their contribution tonet circulation is difficult to quantify, but thealongshore component of winds is highly cor-related with both onshore and alongshore com-ponents of surface and subsurface currents.

Sea level changes on either side of UnimakPass due to storm track and pressure cell move-ment are probably responsible for the fluctua-tion of magnitude and direction in the flowthrough the pass, which at times is southward.These reversals are more likely to occur whenthe flow from the seasonally variable AlaskanCoastal Current, from the Gulf of Alaska, is at itsminimum. The shoaling bottom through UnimakPass gives rise to vertical turbulence and mixesthe water column.

On the north Aleutian shelf, the net north-easterly flow of approximately l-5 cm/s is pres-ent within the coastal zone (Baker 1983; Clineet al. 1982; Thorsteinson 1984). This current isbelieved to be continuous with a weak currentpast Nunivak Island (Kinder and Schumacher1981). Near Port Moller, currents have smallermagnitudes and do not intensify near the coast.Close inshore, within 50 km, currents rangesfrom 1 to 6 cm/s (Kinder and Schumacher 1982).

A weak mean flow shows a cyclonic tend-ency around the perimeter of Bristol Bay, withmaximum speeds (roughly 3.5 cm/s) found nearand inside the 50-m isobath and in the coastaldomain. Mean speeds observed in the centralshelf domain were less than 1.0 cm/s, with nosense of an organized circulation (Kinder andSchumacher 1981). There is apparently a netwestward convection of water from the centralbasin of Bristol Bay into the Bering Sea.However, flow in this central region is highlyvariable, atmospherically forced, and difficult to

quantify. Coastal waters along the northernboundary of Bristol Bay, also called the coastalcurrent, continue to follow the bathymetry. Thecoastal current flows northwesterly into the Ber-i n g S e a a n d t h e n n o r t h e r l y a l o n g t h eYukon/Kuskokwim Delta.Thus, the fundamentalcirculation in outer Bristol Bay consists of atypically unclosed, counterclockwise gyre opento the Bering Sea and driven by acombination ofwind, tide, estuarian, and thermohaline effects.

Ninety to ninety-five percent of the velocityvariance within the bay is tidal, with tidal cur-rentsan order of magnitude largerthan the meanflow. For example, on the north Aleutian shelf,where net currents are only l-5 cm/s and thetypical wind-driven currents are approximately10 cm/s at 5 m, the tidal currents are 40-80 cm/sor more (Thorsteinson 1984). Turbulence result-ing from tidal currents causes mixing of thewater column from the bottom to about 50-mabove the bottom. Tidal currents in Bristol Bayare nearly reversing along the Alaska Peninsulaand become more cyclonic and rotary offshore.National Ocean Survey current tables show achange in maximum ebb currents from 20-25cm/s up to 30-40 cm/s in June near Amak Island.Near Port Moller, the tidal current speeds are ashigh as 100 cm/s (U.S. Department of Commerce1980). At a depth of 2m the calculated tidalresidual current is approximately 3-4cm/s,spatially highly variable, and directed to the nor-thwest (Leendertse and Liu 1981).

Kinder and Schumacher (1981) identifiedthree separate hydrographic flow regimes in thesoutheastern Bering Sea. The Coastal regime ispresent inside the 50-m isobath in the vicinity ofNunivak Island. It is characterized by generallywarm, low saline, vertically well-mixed waterwhich has typical currents on the order of2-5 cm/s toward the northwest. The Middleregime is present in the central Bristol Bayregion, where water depths are on the order of50 to 100 m. It isdividedfrom thecoastal regimebyafront with an enhanced salinity gradient andis characterized by a strongly stratified, two-layered structure extending approximately tothe 100-m isobath. Mean flow is generally lessthan 1 cm/s, with no characteristic vector-meandirection. The Outer hydrographic region isdivided from the middle region by a front alongthe 100-m isobath and is present out to the shelfbreak in the open waters beyond Bristol Bay. Afine vertical structure separates surface layersfrom the deeper, more well-mixed layers. Thevector-mean current in this regime is directed to

the northwest, with magnitudes on the order ofl-10 cm/s and a statistically significant cross-shelf component of about l-5 cm/s.

Yukon Delta

The dominant current near the Yukon Deltais the northward flowing Alaskan Coastal Water.The current is thought to bifurcate at the north-west cornerof the delta, with one fork flowing in-land, toward Norton Sound, and the remainingflow continuing northward (U.S. Navy 1958).Local and seasonal effects can produce variabil-ity ih the prevailing flow directions. In winter,when winds are from the north, flow offshore ofthe delta can actually reverse for days or weeksat a time (Aagaard and Coachman 1981). Thissituation accounted for the flow of the AlaskanCoastal Water about 30% of the time betweenSeptember 1976 and March 1977 (Zimmerman1982). The surface currents offshore of the deltatend to flow in the same general direction as thesynoptic and mesoscale winds, from the north ornortheast in winter and from the southwest dur-ing open waterseason.The typical summerwindfrequently produces downwelling and shore-ward transport of water, which results in a raisedwater level and increased wave energy near thecoast.

Norton Sound

The currents in Norton Sound are domi-nated by regional wind and surface pressure pat-terns. The highest observed flow was measuredat about 50 cm/s; flow decreased with increas-ing depth (Muench 1981). These atmosphere-driven flow events may differ from the mean flowand produce uncertain, intermittent variability inthe circulation pattern. Oceanographic datafrom the mouth of Norton Sound indicate a netnorthward water transport, with strong seasonaldifferences in movement rates. Currentsb e t w e e n t h e m o u t h o f t h e s o u n d a n dSt. Lawrence Island to the west are characteriz-ed by somewhat pulsive north-south flow eventshavingspeedsof50-100 cm/s(Muench, Pearson,and Tripp 1978). These speeds contrast with re-ported mean flow rates of 15 cm/s observed inrelative synchrony with major meteorologicalevents. The mean circulation pattern within thesound iscyclonic incharacter(Drakeet al. 1980).A typical feature is westerly flow of water mass,varying in extent and intensity over time, alongthe northern coast (Cline, Muench, and Tripp1981). The tidal component in the sound ison the

I=6

order of 50 cm/s and reverses either diurnally orsemidiurnally. The reversals are roughly north-east/southwest within Norton Sound.

The upper- and lower-layer circulation isdecoupled in the eastern sound, but less so inthe western sound, where there is a monotonicdecrease in speed along with a slight rotation offlow as depth increases. Northwesterly surfaceflow rotates to westerly near the bottom. In sum-mer, easterly flow enters the sound along itssouthern shore, curves cyclonically to the north,and is then deflected to the west at the north

coast, roughly following the bathymetry. Thisflow varies in intensity and extent from year toyear. In the summer of 1979, a westerly meanflow paralleled the coastline and was super-imposed upon a highly variable flow whichincluded reversals (Muench 1981).

Bering Strait

The Bering Sea is characterized by an openshelf south of St. Lawrence Island. Mean cur-rents are variable in direction and range from 1 to

4 cm/s, with the tidal current accounting for55 (f 31)O/0 of the fluctuation (Coachman, Salo,and Schumacher 1983). Near St. LawrenceIsland, the Bering Sea narrows into two straits,the Shpanberg and Anadyr. North of the islandthe two straits merge to form the Bering Strait.Circulation here is dominated by a northwardmean flow ranging from 4 to 15 cm/s, with verysmall tidal influences 24 (+- 13)Y0 variability(Coachman, Salo, Schumacher 1983). Flow inboth the Anadyr and Shpanberg is to the north,approximately parallel to the local bathymetry.The flow appears to come from around both ends

i/\

+ \// /

Figure 4. Bering Sea Currents-Summer

l-7LegendBering Sea surface currents. Numbersindicate mean speed in cm/s. Arrowsdepict flow as follows:

h Prevailing current direction

e - - Variable current direction

Bering Sea surface currents synthesized fromArsen’er 1967; Goodman et al. 1942; Kinder andSchumacher 1981; LaBelle 1983; Marine AdvisoryProgram, University of Alaska: Notorov 1963;Pelto 1981; Takenouti and Ghtahi 1974; and U.S.Navy 1977.

of St. Lawrence Island. Frequent reversals arecoincidental with meteorological events. Thesereversals can affect the flow over vast regionscovering thousands of square kilometers. Thepresence of ice appears to dampen the impact ofwind stress forcing. The major driving force forthe northward flow through Bering Strait is thesea surface sloping down to the north (Aagaardand Coachman 1966). A slope of 2 x lo- 6 isassociated with average summer northerlytransport of approximately 1.6 x10” m3/s. The nor-mal condition is, thus, one in which sea level inthe southern Chukchi Sea (in summer) is about

1 6 0 ” l650170” n 187

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0.5 m lower than in the northern Bering Sea. A ma-jor cause of variations in the sea level differencemust lie in fluctuations of the regional winddistribution. It is also possible that the at-mospheric pressure field may itself directlymodify the oceanic pressure field (Aagaard,Coachman, and Tripp 1975).

An examination of recent meteorologic data(Aagaard and Coachman 1981) showed thefollowing results. In every case of southerly flowthrough the Bering Strait, the large-scale atmos-pheric pressure patterns were the same. One day

\

Figure 5. Bering Sea Currents-Winter

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before a peak in southerly flow, a strong low-pressure system was centered some distance tothe southeast of Bering Strait in the area ofBristol Bay, Kodiak, Anchorage, and the northernGulf of Alaska. At the same time the Siberian highwas centered some distance west or west-northwest of the strait. The isobars signifying thestrongest pressure gradient between pressurecenters were located precisely over the BeringStrait region. Most significantly, they had a nearlynorth-south orientation which extended fromover the Chukchi Sea south into the central Ber-ing Sea-completely across the northern BeringSea shelf. If the north-south orientation of theisobars did not extend totally across the northernshelf or if the isobars were oriented northeast-southwest (the nearest typical configuration),strong southerly flow events did not occur.

The mechanism which drives major southflow events now seems clear. Strong north windsmust develop over the entire northern Bering Sea,not just over the immediate region of BeringStrait. Large-scale, strong atmospheric pressurecells are required: a low far to the southeast and ahigh well to the west. The strong northerly windsgenerated thereby move water southward off theentire northern Bering Sea shelf. Removal of suf-ficient water off the northern shelf generates asea-level slope down to the south-sea-levelslope has been shown to be the major forcedriving transport through the strait (Coachmanet al. 1975). This, together with the strong northwinds caused by the east-west atmosphericpressure gradient, drives enhanced southerlytransport. These conditions apparently requireabout one day to develop, so that maximum southtransport occurs the following day. Because the

system behaves to a marked degree as a coherentunit, water levels at both St. Lawrence Island andCape Lisburne fall together and are nearly inphase with the transport.

Northward transport stands in contrast tothe southerly transport events. Periods of north-erly flow tend to be more persistent and not sogreat in magnitude, nor do they show the markedepisodic character of the southerly flows. Thegreater persistence of northerly flow must reflectthe basic driving force, a higher sea level in theBering Sea than in the Arctic Ocean (Coachmanet al. 1975), which still remains unexplained.There were, however, a number of relatively rapidnorthward accelerations of transport during theseven months of record which appear to have twobasic causes:

(1) After strong south transport events,rapid accelerations commonly occur which canbe thought of as compensatory. When atmos-pheric conditions causing the southerly trans-port event dissipate, water is not being removedfrom the northern Bering shelf, but there is stillvoluminous southerly transport in the system.Water “p i les up” in the region aroundSt. Lawrence Island and Norton Sound, a condi-tion reflected by a strong, positive difference inwater level. Following this by about one day, astrong northward acceleration occurs.

(2) Occasionally, major northward accel-erations appear to be, at least in part, directlydriven by atmospheric conditions. Specifically,these are a strong low pressure centered in thewestern Bering Sea southwest of Bering Strait,or a deep trough from the central Aleutians

toward the northwest, so that the isobars in thestrong pressure gradient are directed northwardfrom the central Bering Sea along the axis of thesystem. This configuration creates strong,southerly winds which can move water from thecentral Bering Sea onto the northern Bering Seashelf, raising the water level in the vicinity ofSt. Lawrence Island and enhancing the sea-levelslope down to the north.

Central Bering

West and northwest of the North AleutianBasin and Yukon Delta lies St. George Basin,the Central Bering Sea, and still further west, theNavarin Basin. Circulation in these regions isnot as well understood as in the coastal basins.Fewer studies have been conducted in the off-shore Bering. Data are site-specific andsporadic over decades. No consistent flow pat-terns have emerged as representative of theregional circulation. In fact, there is little con-sensus among investigators that the principalflow is north-south, east-west, or cyclonic oranticyclonic in nature (See Natarov 1963;Arsen’ev 1967; Tak enovti and Ohtani 1974;Goodman 1942; Ratmanov 1937). The northwardflowing, eastern boundary current is roughlybalanced by a southward flow along the Sovietcoast. Within the central region, flow is probablydominated by the location and strength of large-scale atmospheric pressure cells. Responsetimes, directions, and persistence are probablyof a similar scale as those controlling flowthrough the Bering Strait (Aagaard andCoachman 1981). Thus, a dominant regionalflow pattern is not readily observed nor easilyquantified.

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Sea Ice

Introduction

The annual cycle of formation and dissipa-tion of sea ice in Alaska waters has widespreadeffects on a number of phenomena. When the iceforms, the coastal climate changes in characterfrom maritime to continental with much coldertemperatures and lower humidities than wouldbe the case if open water were present. The icealso interferes and even stops water transporta-tion with the possible exception of icebreakersand otherspeciallydesigned ships. It makes thecleanup of oil spills difficult, if not impossible,by hampering the operation of cleanup equip-ment and by trapping oil under the ice. Sea icealso has important effects on the life cycles ofliving creatures in and near the sea.

In the Bering Sea, the sea ice generallybegins as fast ice formation along the shores ofthe Seward and Chukotsk peninsulas inOctober. As the season progresses and watersin the more open portions of the Bering Sea cooloff, the pack ice generally begins its seasonalsouthward formation in November. An esti-mated 97% of the ice in the Bering Sea is formedwithin the Bering Sea(Leanov 1960); very little istransported south through the Bering Strait.During periods of increasing ice and prevailingnortherly winds, the ice apparently is generatedalong the south-facing coasts of the Bering Seaand moves southward with the wind at as muchas 1 knot before melting at its southern limit(Pease 1971). During periods of southerly winds,ice coverage generally decreases in the Bering.Prevailing winds can persist in one direction forweeks at a time in winter in the Bering Sea,causing a wide variation in ice cover from monthto month and from year to year(see Figure6 andmap set 17, Section II of this volume).

Recurring Leads and Polynyas

Wind and current stresses on the ice cancause tension or divergence and open relativelynarrow, long stretches of open water in an other-wise dense ice cover. In the absence of strongcurrents, the wind induces leads which runperpendicular to the wind direction. Flaw leadsgenerally occur just seaward of the stable fastice zone when strong offshore winds develop.

In the Bering Sea a wind-induced polynya(Figure 6) immediately south of St. LawrenceIsland is a frequent but undependable feature(McNutt 1981; Wohl pers. comm.). Northerly

‘ 6 4 ”

G a m b e l l ~

172",

n a l a k l e e t 64’

/,f%t. Michael

held&

Figure 6. Recurring Polynyas

winds cause the polynya to form in the leeof theisland as sea ice is advected to the south. Thepolynya can extend more than 160 km and is fre-quently covered with thin ice. However, thefeature is temporal, and a wind shift to southerlyflow can close this area rapidly. At such times, ac o r r e s p o n d i n g p o l y n y a t o t h e n o r t h o fSt. Lawrence Island is sometimes observed, butit is generally much smaller and occurs less fre-quently.

A polynya can form on any side of NunivakIsland, depending upon the prevailing winddirection. Usually the feature is located to thenorth or south, under southerly or northerlywinds, respectively. Like the polynya off

Synthesized from: McNutt 1981; Stringer, Barrett, andSchreurs 1980; Wohl 1982.

St. Lawrence Island, the appearance of thispolynya is variable, but it is usually observed atleast once each year, often more. Its extent isvariable, and thin ice commonly covers thepolynya quickly during cold, northerly windstorms.

The polynyas shown for Norton Sound,Bristol Bay, and Kuskokwim Bay were takenfrom Stringer, Barrett, and Schreurs (1980).These features were mapped from LANDSATscenes collected between 1973 and 1976.Generally, the major polynyas in these areasopen in response to northerly winds, whichcause all but landfast ice to move toward thesouth. As the polynyas are opened by the wind,

l-10

new ice forms and is, in turn, advected south-ward. This mechanism for new ice productioncan be very efficient under the proper cir-cumstances (Pease 1980). None of the polynyascan be considered even semipermanent since areversal in the wind direction can completelyclose them. Furthermore, many of the areasshown are partially covered with very thin icewhen the northerly winds bring below-freezingtemperatures.

Fast Ice and Shear Zones

According to World Meteorological Organi-zation sea ice nomenclature, fast ice includes allice that has become attached to the shore, evenmultiyear pack ice. A common feature at theseaward boundary of the fast ice is an area ofshear ridges. Shear ridges in the Bering Sea tendto be more localized and of lesser extent andmagnitude than farther north. Figure 7 showsthe various kinds of ice near shore. Bering Seaice does not have any multiyear ridges. The ac-companying fast ice boundary maps, from theAlaska Marine Ice Atlas, were synthesized fromStringer, Barrett, and Schreurs (1980).

Any significance accorded to trends appar-ent on these maps must be tempered by consid-eration of the variability exhibited in theice-edge data. At some locations, the edge of thefast ice varies considerably in position duringeach period. Although the average edges alongthe coast show a temporal trend, it has onlyminor significance. In other locations, thevariability of the fast-ice edge of each period issmall compared to the changes in the averageposition from period to period (Stringer 1981).The intraseason and interseason variability ofthe fast-ice edge are very dependent on themeteorology and associated wind patterns aswell as the offshore bathymetry. Although theprevailing winds in winter are generally north-easterly, there are often periods of a week ormore with southerly winds. During northeasterlywinds, shore leads and polynyas open up,only tobe closed again when winds shift to south orsouthwest. Also, the high tide ranges in BristolBay and along the coast south of Norton Soundtend to break up extensive areas of fast ice,except where it is grounded on mud flats oroffshore shoals.

‘330

53"

fNI

62:

..*. ..

/

. .. /

Legend

. . . . . , . . . . 18-21 January, 1975; 1976-22-25 January, 1977- - - - - - - 21-23 January, 1978- - - 20-23 January, 1979

175" 170" @- 165" 160"I 0 , I

_ __ _~ _ _--- _---Figure 8. January Southern Ice Limit for 19751979. source: Niebaur 1981.

Floating Driftingra5L lbci Extension Pack Ice

Ice _/1 IA ,Ub L”ll”T I -

t Active Tidal Cracks

c Traces of Former Active Tidal Cracks

Figure 7. Sea Ice Zones and Types

l-11

Legend---. - Average

Maximum Deviation

05o”KILOMETERS

0 5 0 1 0 0

NAUTICAL MILES

Synthesized from: Stringer, Barrett and Schreurs 1980.

I I I174'174' 170" 166" 164" 162' 160"

/ 1 1FEBRUARY=MARCH

3ERf.Wc SEA

172"1

Figure 9. Seasonal Fast Ice Boundary- Norton Sound (February/March)

l-12I I 1 I

174" 1 7 0 " 166" 164" 162" 160"

- I Yk-

.

BERING SEA

NOR TON SOUND

St. Lawrence lrland

1 7 2 " 1 7 0 " 1 6 8 "

I IFigure 10. Seasonal Fast Ice Boundary-Norton Sound (April/May)

\1 6 6

[MAY-JUNEI\Y++J-;

BERlNG SEA

St. Lawrence Island

1 7 2 "I

NOR TON S

1 7 0 " 1 6 6 "

Kotlik

Y--inflEd”

I&YUKON DELTA “*”

Y

[160'

Figure 11. Seasonal Fast Ice Boundary-Norton Sound (May/June)

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Legend-.-.- Average

Maximum Deviation

050KILOMETERS

0 5 0 1 0 0

NAUTICAL MILES

Synthesized from: Stringer, Barrett and Schreurs 1980.

1 FEBRUARY-MARCH 1

Figure 12. Seasonal Fast Ice Boundary -Southeast Bering (February/March)

Figure 13. Seasonal Fast Ice Boundary- Southeast Bering (April/May)

Bet he

L

Figure 14. Seasonal Fast Ice Boundary-Southeast Bering (May/June)

Tidesl-15

The practical study of tides, aimed at pre-dicting surface elevations and times, involvesthe empirical treatment of observations made atthe desired location over an extended period oftime. The motion of the heavenly bodies, par-ticularly the sun and the moon, relative to theearth is known with great precision, so the tidegenerating potential at any place and time canbe computed. Mathematically the potential canbe resolved into a finite number of strictlyperiodic components which, upon addition, pro-duce the total potential of hundreds of tide-generating components that are listed byvarious authors. Many of the components are ofinsignificant amplitude and can be excludedfrom consideration. In practice only seven com-ponents are widely used: four semidiurnal (M2,S,, N,, K,) and three diurnal components (K,, O,,P,) (McLellan 1965). The names and relativeweights of these components are shown inFigure 15.

Theoretical modelsof tides must beverifiedon the strength of observations at tidal stationswhere the tide wave has been distorted throughpassage over a continental shelf of complextopography. The nature of tides in a particulararea is highly dependent on the bathymetry,shape, and direction of the coastline andlatitude. In the Bering Sea in the last 10 yearsthere have been a large numberof presuregaugeand current meter observations taken so that

tides can be modeled with a high degree ofacuracy. The following discussion paraphrasesmaterial contained in Pearson, Mofjeld andTripp, 1981 in which theoretical tide modelresults are compared to observations. The tidesmost concerned with are the principal tidal con-stituentsN2and M2in thesemi-diurnal band and0 and K in the diurnal band. Ordinarily the S2would be included in thediscussion, however,S$is anonymously small throughout the BeringSea, possibly because it has small amplitudes inthe adjacent North Pacific Ocean. The com-plicated distributions of semi-diurnal and diur-nal tides in the Bering Sea produce a rich varietyof tidal types, ranging from fully semi-diurnal insome regions to fully diurnal in others.

The tide wave enters the Bering Sea as aprogressive wave from the North Pacific Ocean,mainly through the central and westernpassages of the Aleutian-Komandorski Islands.The Arctic Ocean is a minor secondary source oftides which propagete southward into the northBering Sea where they complicate the tidaldistributions.

Tides in the Bering Seaare considered to bethe result of cooscilation with large oceans.Once inside the Bering Sea, each tidal consti-tuent propagates as a free wave subject to Cor-iolis effect and bottom friction.

NAME OF COEFFICIENTSYMBOL PARTIAL TIDE R A T I O

Y Principal Lunar 100.0

sz Principal Solar 46.6

N2 Larger Lunar Alliptic 19.2

K2 Luni-Solar Semi-Diurnal 12.7

K> Luni-Solar Diurnal 58.4

0, Principal Lunar Diurnal 41.5

p, Principal Solar Diurnal 19.4

Contracted from table 15.1,Elements of Oceanography (McLellan 1965).

Figure 15. Major Tide Components

The tide wave propagates rapidly acrossthe deep western basin. Part of it then pro-pagates onto the southeast Bering shelf wherelarge amplitudes are found along the AlaskaPeninsula and in Kvichak and Kuskokwim Bays(Figure 17). Another part propagates north-eastward past St. Lawrence Island and into Nor-ton Sound. Over most of the Eastern Bering Shelfregion the tide is mainly semi-diurnal, but in Nor-ton Sound diurnal tides predominate. Over theremainder of the Bering tides tend to be mixed.In the Aleutians diurnal rather than semi-diurnalcomponents are stronger.

Legend

Diurnal range is the average difference in height between mean Diurnalhigher high water and mean lower low water in feet on a single Rangeday. Max M i n

Max diurnal and Min diuranl are the maximum and minimum dif-diurnal diurnal

ferences in feet respectively between the higher high water and

=F

Max M i nlower low water that are predicted to occur during the year. tide tide

Max tide is the highest tide predicted to occur at the location in feet above the datumlevel generally taken to be the mean of the lower of the two low waters of each day.

Min tide is the lowest tide predicted to occur at the location in feet above the datumlevel generally taken to be the mean of the lower of the two waters each day. Anegative number indicates a level below the datum level.

Prepared by AEIDC from West Coast of North and South America Tide Tables, Highand tow Water Predictions, 1986, West Coast of North and South America,NOSINOAA, 1985.

l-16

6 KashegaBay

10.817.8 5.9*13.4 -4.8

11 PortMoller

17.310.4 2.5*9.3 -2.5

12 SandPoint

12.319.9 6.6*14.9 -5.2

18.226.1 11.0*22.1 -4.6

22.631.6 11.0*26.9 -4.7

15 Naknek A.Entrance

~ 13 Port~ Heiden II 14 Egegik R.

Entrance

3.2

I 16 NaknekAir Base

16.5

17 Kvichak

19.528.4 12.0*24.0 -5.0

16 NushagakBav

8.511.9 2.6*11.3 -0.6

I 19 GoodnewsBav I

12.217.3 4.4*16.8 -0.5

I 20 KuskowakCreek

4.01.2 0.3*1.0 -0.0

21 Bethel

I24 ;N;;east 1 125 Mei:oweyiki

26 CapeRomanzof

2.32.5 1.0*2.1 -0.3

27 KwikluakPass

26 Apoon 29 St.Mbuth 1 1 Michael 1

1.83.1 0.4*2.6 -0.5

30 CarolynIsland

1.61.6 0.4*2.1 -0.6

31 Nome

Figure 16. Tide Data

Legend

Type of Tide

pm Semi-diurnal

r=AD i u r n a l

1-1M i x e d

-8- Corange (Feet)

Adapted from Bering, Chukchi, and Beaufort SeasCoastal and Ocean Zones Strategic Assessment DataAtlas, Prepublication Edition, and U.S. Navy, 1972.

160’ ! ’ 170:650 .---c n 18d

L---,/ Khatirka-In-Chwkot A

1x

No

Northeast Ca

\

Figure 17. Tide Co-Range

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Storm SurgesStorm surges are waves oscillating in the

period range of a few minutes to a few days, in acoastal or inland water body, resulting from forc-ing from atmospheric weather systems (Murty1984). By this definition, wind-generated waves(often referred to as wind waves) and swell,which have periods of several seconds, are ex-cluded. The spectrum of storm surge waves iscentered around lo-” cycles per second (CPS),which gives a period of about three hours. How-

Legend

t---Secondary Track

Storm tracks depicted are a com-posite of actual flooding eventsfrom case histories. Wise, et al.,1981.

ever, depending mainlyon the topography of thewater body and secondarily on other param-eters, such as the direction of movement of thestorm, strength of the storm, stratification of thewater body, presence or absence of ice cover,and nature of tidal motion in the water body, theperiods of the water level oscillations may varyconsiderably. Even in the same water body,storm surge records at different locations canexhibit different periods.

Although storm surges belong to the classknown as long waves, as do astronomical tidesand tsunamis, there are at least two importantdifferences. First, whereas tides and tsunamisoccur on an oceanic scale, storm surges aresimply a coastal phenomenon. Second, signifi-cant tides and tsunamis cannot occur in an en-closed, small, coastal or inland water body, butstorm surges can occur even in lakes, or incanals and rivers. The range or height of a storm

Figure 18. Storm Tracks with Storm Surge Floods

I-20

surge depends not only on characteristicsof thestorm but also on the topography onshore andbathymetry offshore. Shallow water bodiesgenerally experience surges with greaterranges. Also, the height of a storm surge is less ifthe sea floor is steep than if there is a shallowslopetotheseafloor(Murty1984).Stormcharac-teristics that effect the height of a surge includeatmospheric pressure; wind speed, direction,and length of fetch; the latitude; and the direc-tion and speed of storm movement. Air andwater temperature differences also affect theheight of surges.

Following is a discussion of the stormsurge potentials from astudydone in 1981 (Wise,Comiskey, and Becker), supplemented by stormstatistics since then (NOAA Storm Data,1981-1986) and a modeling study for surges inNorton Sound (Wise, Comiskey, and Becker1981; Kowalik and Johnson 1985).

Along the southwest and south coasts ofthe Seward Peninsula the terrain is generally ofmoderate relief, with the exception of PortClarence and the east end of Norton Sound. Thewaters offshore are shallow with a gentlysloping sea floor. The open waters of the BeringSea provide a long fetch for the development ofstorm waves. Sea ice restricts the developmentof storm waves from about the first of Decemberto the first of June on the average; however,there can be a high degree of annual variability.Of 13 known flooding events in Nome, all excepttwo occurred in the fall. One destructive stormfor which little factual information is availableoccurred in April 1906, and another occurred inJuly1969.Surgesabove4 m(12 ft) haveoccurredalong this section of coast. The most recent wasin November 1974; a 3 m (lo-ft) surge broughtwater into the town over the sea wall. This par-ticular storm caused widespread flooding allalong the Bering Sea coast.

With the exception of the Shaktoolik Rivermouth, which is of low relief and marshy, thecoast at the east end of Norton Sound isgenerally rugged due to the proximity of theNulato Hills. The south coast of Norton Sound isgenerally of low relief. The sound itself isshallow, with a gently sloping sea floor that isvery favorable for the development of stormsurges. The range of wind directions for thedevelopment of storm surges is limited to west-

southwest to west. However, the east end ofNorton Sound often experiences minor flooding,despite unfavorable winds, due to rising sealevels all over the sound.

A storm surge modeling study(Kowalik andJohnson 1985), determined that for theNovember 1974 storm the highest surge inNorton Sound was in Norton Bay, with amodeled surge of more than 3 m (10 ft). Thesame study also determined that negativesurges of more than a meter can occur in theeastern end of Norton Sound in the presence ofstrong, persistent northeast winds in winterwithan ice cover present. The winds tend to move thepack ice away from the shorefast ice and reduceice cover from 0.7 or 0.9 coverage to less than0.55 coverage over much of Norton Sound.

Eleven of the twelve storm surge cases atUnalakleet occurred in the fall; the other was inJuly. Sea ice and shorefast ice limit the fetch forthe development of positive storm surges fromabout the first of December to the first of June.

The shores of Pastol Bay and the northcoast of the Yukon River delta do not have longfetches favorable for the generation of wavesand storm surges. However, this area ex-periences surges due to increases in the heightof the water in Norton Sound. The remainingcoast of the Yukon Delta is exposed to the openwaters of the Bering Sea, where conditions arevery favorable for the development of stormsurges due to low retief onshore, shallow wateroffshore, and thousands of miles of open sea.Most surges causing property damage occur inthe fall or in August. However, early summersurges can be very hard on nesting birds in thesalt flats. In June 1963 80% to 90% of the blackbrandt production was lost due to flooding of thenesting area after eggs were laid.

The coastal area is generally of low relieffrom the Kuskokwim Delta to Goodnews Bay,with numerous lakes, sloughs, and marshes.From Goodnews Bay to the Nushagak Peninsulathe coastline is more rugged due to the proximityof the Ahklun Mountains. The remainder of thecoastline of Bristol Bay is similar to the stretchfrom the Kuskokwim River to Goodnews Bay. Of-fshore the shape of the sea floor is conducive tothe formation and enhancement of stormsurges. From Goodnews Bay northward an ade-

quate fetch can be generated with storm windsfrom south through west to northwest. East ofGoodnews Bay, westsouthwest to west are theonly directions from which an adequate fetchcan develop.

Autumn and late summer are the seasonsfor destructive storm surge flooding in this area.There are ten known cases of storm surgeflooding of populated areas; seven were inautumn and three were in August. Two storms, inNovember 1979 and in August 1980, account formost of the factual reports of storm surgeflooding. The November 1979 storm causedstorm surge flooding from Cape Newenham toScammon Bay. Surges were estimated at 2.5 m(8 ft) in exposed locations in the KuskokwimDelta. The storm was on a track from west-southwest to eastnortheast, and a long fetch ofmore than 640 km (400 mi) developed with thestorm. The August 1980 storm, one of the fewsummer flooding events, caused flooding on theshore of Bristol Bay. The storm was on a trackfrom south-southwest toward north-northeastfrom near Atka Island, in the Aleutians, toKuskokwim Bay. Exposed locations showedsurge flooding up to 4 m (12 ft).

The coastal area from Hooper Bay to KinakBay is a favored nesting area of migratory birdsin the spring and summer. Minor storm surgesthat cover nests in this area at the wrong timecan be detrimental to the annual production ofseveral species of birds. Five cases of minorsummer flooding of the salt flats were docu-mented in an annual report for the ClarenceRhode National Wildlife Refuge (USFWS 1964).One event (June 22,1963) caused a loss of blackbrandt offspring estimated at 80% togO% of theyear’s production.

Most of the north shore of the Alaska Penin-sula east of Cold Bay is favorable for the occur-rence of storm surge flooding, with low, marshyterrain onshore and a moderately sloping seafloor offshore. West of Cold Bay the AleutianIslands and the south shore of the Alaska Penin-sula conditions are not favorable due to ruggedterrain onshore and steep ocean floor offshore.The only storm surges discovered in this areawere at Meshik, or Port Heiden, and St. Paul.Damage to structures in this area is more likelyto be from strong winds and beach erosioncaused by wave action than from flooding.

Superstructure IcingStructural icing on ships, offshore struc-

tures, and port facilities isawintertime hazard inopen waters and coastal sections of Alaska. Theicing causes slippery decks, renders movingparts inoperable, and, in extreme cases, causesuneven loading and raises the center of gravityon small ships. Accumulation of ice on riggingand on deck equipment such as crab pots alsoincreases wind effects because a larger surfacearea is presented to the wind. Ice forming onstructural surfaces above or close to a body ofwater arises principally from seaspray(Naumanand Tyage 1985; Liljestrom 1985), with lesseramounts from atmospheric precipitation (freez-ing rain and wet snow) and fog (arctic sea smoke,white frost, black frost). Sea spray, the mostdangerous source of icing, is produced by thebreaking of waves against obstacles such asships’ hulls, other floating objects, shore struc-tures, and, possibly, other sources(Minsk 1977).

Statistical analysis (Borisenkov and Panov1972) of more than 3,000 cases of ship icingindicates that in 86% of the cases icing wascaused by ocean spray alone. Spray combinedwith fog, rain, or drizzle (l iquid sources)accounted for only 6.4% of the cases, and spraycombined with (solid source) snow only 1.1%.The cases of icing attributable only to fog, rain,or drizzle account for 2.7% (Minsk 1977). In theremainder of icing cases data were not suffi-cient to determine the cause.

Since the overwhelming majority of super-structure icing on ships and offshore structuresis from sea spray, the remainder of this sectionwill concentrate on this type of icing. Since aship can present different aspects to the windand spray, it is to be expected that the amount ofspray reaching the ship will vary: Russianobservations (Kultashev, et al. 1972) showedthat the greatest frequency of spray and, there-fore, icing occurs when aship is heading into thewind at an angle between 15” and 45”. Asym-metrical icing occurs under this condition, withthe greater accumulation on the windward side.Less icing occurs with the ship headed directlyinto the wind, and then accumulation tends to beuniform. With ships heading downwind, sprayicing is generally much less than at other angles.In developing the nomogram for forecastingspray icing potential, downwind cases(thoseforwhich the ship’s heading was 120”orgreateroffthe wind) were not used.

Meteorological/oceanographic conditionsnecessary for significant spray icing are watertemperatures less than 8X, winds of 25 knots(13 meters per second) or more, and airtemperatures less than - 2 “C (28”F, the freez-ing temperature of seawater of average salinity).Generally, the stronger the wind, and the colderthe air and water, the higher the rate of icing oncomparable vessels or structures. In some

Figure 19. Superstructure Icing Rate Nomogram

cases, however, where the wind fetch is notsufficient to fully develop waves, icing rates arelower.

The accompanying potential superstruc-ture icing rate nomogram (Figure 19) is amodification of that shown in Wise and Com-iskey(1980), using theopenoceancasesappear-ing in Pease and Comiskey (1985), developed

3 23 0

2 8

2 6

2 4

2 2

c 2 0

; 18

3 16

3 14m=E 12

Q)l- 10

4

2

0- 2

-4-6

-8

Wind (Knots) Sea Temperature (F)8 0 7 0 6 0 5 0 4 5 4 0 3 5 3 0 2 5 OcuOwwON-rwmmmmm*~-r*

hllI\I\\!\‘4\\ r\ \-

Instructions: The nomogram is enteredwith air temperature and wind data. Atintersections of these proceed diagonal-ly to sea surface temperature and readicing classification (light, moderate.etc.). The related range of icing is shownin the table below.

Icing Classification Rate per Hour

Adapted from Wise and Comiskey, 1980, using data fromPease and Comiskey, 1985.

from icing case histories in the Gulf of Alaskaand southern Bering Sea. Icing intensities in in-ches per hour are also from Pease and Comiskey(1985). If a vessel experiencing icing takesevasive action (i.e., changes heading, reducesspeed, seeks shelter, etc.), icing rates experienc-ed would probably be less.

Reported cases of ship icing (Figure 20) inthe northern Gulf of Alaska and the Bering Seaare shown from two sources; Borisenkov andPanov (1972) and WBH29 (Dyson 197583). The

lackof reported icing in the northern Bering maybe a result of reduced ship traffic as well as con-ditions not favorable for icing. Kozo (1983) esti-mates a potential for superstructure icing for thenorthern Bering Sea in September, extendinginto Norton Sound in October, and even into thesouthern Bering during the most extreme condi-tions. Icing potential decreases in the northBering as the sea ice cover advances; however,the potential for moderate or heavy icing down-wind of the sea ice edge persists throughout thewinter.

Legend

Superstructure icing reports9 received by NBH29

0 Approximate locations oficing reports from V.V. Panov

Figure 20. Related Occurrences of Superstructure Icing on Ships

l-23

Figure 21. Climatic Means and Extremes

2 5 . 2 4 15.1610.50) 2.77 5.45 ) 2.80

6 9 . 4 5 6 . 968.7 1 12.5 23.0 1 NA

llfi AR

18.959.56 12.06

75.339.2 1 20.6

1 0 8

v

14.16 19.758.07 Il.53 5.65 12.00

39.9 63.737.3 I 14.0 27.5 1 NA

14.777.82 12.99

55.825.9 1 9.0

7 4

16.8713.68 16.60

67.2

18.615.86 12.01

98.148.6 1 14.1

.-

q%

4 Moses Point 5 Unalakleet 6 Emmonak C a p e R o m a n z o f 8 NunivakG a m b e l l 2 Northeast Cape

12 King Salmon 13 St. Paul Island 14 Port Heiden

43.51N A 1 5 . 0 0

92.1NA 1 10.4

?C

3 6 . 4 211.421 3.13

20.7 [ 13.15 1

-7E-i-s19.33

7.30 12.0043 7

15.685.55 13.00

4 9 . 43 7 . 7 18.0

-+

-*16.13

5.81 12.024 6 . 8

21.6 1 8.2

5 1 . 5 713.391 3.90

4 0 . 523.3 1 9.0

<A20.0.~~- 11.0

-*

JI11 Dillingham

1 , I5 Port Maller 16 Sand Point9 Bethel

I1 7 C o l d B a y

IO Cape Newenham

16 0 . 5 3

16.041 7.59I72.252.6 ) NA

18 Dutch Harbor

-7%--k-3 0 . 4 4

8.71 15.206 9 . 6

42.1 1 NA

-*23.41

NA 12.503 3 . 5

NA 1 10.4

-*

6 1 . 1 917.34 5.49

9 9 . 254.8 1 17.0

2 0 . 5 57.63 13.08

70.240.1 ) 30.0

*

I Driftwood Bay 0 Nikolski 1 Adak

Total Precipitation <6 1 . 1 9

17.34 I 5.49

Snowfal l -* 54.8ggi2, 7.0

2 Shemya

Temperature(°C)

Mean Annual Maximum Mean Annual Minimum

Highest Lowest

-

Total Precipitation (inches)(Rain and Snow)

Average Annual

Greatest Month Greatest Day

-

Snowfall ( inches)

Average Annual

Greatest Month Greatest Day

-

Snow depth (inches)

Annual Maximum on Ground-

Surface Wind (knots)

Prevailing Direction Average Annual Speed

Fastest Direction and Speed

17.6 1 3.6

22.2 1 -14.4

1 1 4 . 3 0

3 4 . 8 7I

7.59

-

58.5

55”

+i

h

4 0-

SSE /9.3

SSE/ 52

Where the mean annual temperature (average of the meanannual maximum and minimum) is less than zero sometype of permafrost will probably be present.

These data can be used for long-range planning anddesign criteria. More detailed information can be obtainedfrom the National Climatic Data Center and the ArcticEnvironmental Information and Data Center in Anchorage.

NA = Information is not available

Prepared from NOAA/NESDIS and Canadian AES data.170” I 180’

l-25

Hypothermia is the cooling of the body’score temperature to 95°F or below. It can causeshivering, numbness, and disorientation. In theextreme it can cause death. The body loses heatgradually in cold, dry conditions, but quicklybecomes hypothermic in wet conditions. Rain,immersion in cold water, and perspiration can allcause rapid heat loss. However, the evaluationand treatment of hypothermia, whether wet ordry, on land or water, is essentially the same,namely to warm the victim by whatever appro-priate means are available.

The following discussion was taken in partfrom Peters (1982).

The body loses heat in five ways:

l A large amount of heat is lost from the bodyin respiration. Exhaled warm air is replaced bycooler inhaled air, producing a net heat loss.Theamount of the net heat loss can be reduced bycovering the mouth/nose with wool or fur, there-by “prewarming” the inhaled air as it passesthrough the material which has been warmed byexhaled air and by heat radiating from the body.

l Evaporation of perspiration from the skinand moisture from the lungs contributes greatlyto the amount of heat lost by the body. Althoughevaporation cannot be prevented, the amount ofevaporation (and therefore cooling) can be con-trolled. Wearing clothing that can be opened orremoved easily for ventilation will let watervapor escape and not condense to liquid water inthe clothing. Keeping clothing dry preserves itsinsulating value and reduces heat loss.

l Sitting on snow, touching cold equipment,and being rained upon are all examples of howheat can be lost as a result of conduction. If anindividual becomes wet a tremendous amountof body heat is lost rapidly. Deaths haveoccurred as a result of immersion in water below40”F-body temperature could not be main-tained. Although not as immediately serious,perspiration, rain, or wet snow should never beallowed to saturate articles of clothing, as thisseriously reduces their insulating properties.

l Radiation causes the greatest amount ofheat loss from the body from uncovered sur-faces, particularly the head, neck, and hands.Coverage of these areas, therefore, is extremelyimportant in keeping warm.

e The body continually warms (by conduction)a thin layer of air next to the skin. If the warmlayer is removed by wind or air currents (advec-tion), the body is cooled. The primary function ofclothing is to retain this layer of warm air next tothe skin by enclosing air in cell walls or betweennumerous fibers, while allowing water vapor topass outward. Heat is lost rapidly with thelightest breeze unless the proper type ofclothing is worn to prevent the warm air from be-ing advected away.

Deaths have been attributed to a loss ofbody heat at temperatures of 4O”F, with a30 mph breeze. Undertheseconditions, thecool-ing effect on the skin is equal to that of muchlower temperatures due to increased evapora-tion and convection. With lower temperaturesand/or strong winds, cooling occurs even morerapidly. Wind protection and insulation (dead airspace) can help ensure that body heat isretained at a safe level.

TreatmentRecognition and proper treatment of hypo-

thermia must be prompt. Delays even afterrescue can cost a person his life. Low bodytemperature is the best indication of hypo-thermia. Blood pressure and pulse are alsogoodindicators. The pulse is generally slow andirregular, while blood pressure is low.

The hypothermia victim is pale in appear-ance, the pupils are constricted and react poorlyto light, and respiration is slow and labored. Hewill usually be shivering violently, with frequentmuscular rigidity. There may also be anappearance of intoxication.

Emergency treatment must begin as soonas possible to stop the drop in body temperature.Wet clothing should be removed. If the bodytemperature is97”Forabove, no treatmentother

than dry clothing and moving the victim to awarm area is generally necessary. If these arenot available, the wet clothing should not beremoved.

Combatting “afterdrop” in the core bodytemperature is extremely important. When heatis applied to the arms and legs, it causes thoseblood vessels to relax. This allows cold blood toflow back into the body core, further cooling thevital organs. Warming of the trunk of the bodyshould be the prime concern.

During experiments in conjunction with theU.S. Coast Guard, researchers determined thatthe best warming technique was from the insideout, by having the victim breathe moist, warmedoxygen (Wilson 1976).

The next best treatment is a hot bath, withthe water temperature between 90 and 100°F. Ifa tub is not available, an inflated life raft could beused. If possible, the limbs should remain out ofthe water. When no tub-type facility is available,a hot (115°F) shower while wrapped in towels orblankets is preferable.

When hot water for a tub or shower isunavailable, wrap the victim in blankets in awarm room with a heating pad or well-wrappedhot water bottle on the chest, or apply bodywarmth by direct contact with a rescuer.

Warm liquids may be given, but care mustbe taken to insure the victim is conscious anddoes not breathe the liquid into his lungs.Alcohol should never be given because it causes“afterdrop.” Observe the victim’s respirationclosely and monitor for vomiting.

It has been learned in studies done inAlaska that victims of wet hypothermia can sur-vive for a prolonged time in cases of deep cool-ing. Apparently, in the rapid cooling whichoccurs with wet hypothermia, physiologicalchanges undergone by the body are more likelyto be reversible than in the slower cooling of dryhypothermia. There have been victims of immer-sion hypothermia who were apparently dead butrevived with proper treatment.

I-27

Wind Chill (Equivalent Temperatures)

The temperature of the air is not always areliable indicator of how cold a person will feeloutdoors. Other weather elements, such as windspeed, relative humidity, and sunshine (solarradiation), also exert an influence. In addition,the typeof clothing worn, togetherwith the stateof health and the metabolism of an individual,influence how cold a person will feel. Coolingmay be described as loss of heat from exposedflesh. Freezing occurs when there is such totalheat loss that ice forms in the exposed tissues.Thecooling poweroftheatmosphere(bywind)isprimarily heat transfer by advection-in humancases, by exposure of uncovered flesh to theenvironment. Even small amounts of air move-ment have considerable chilling effect becausethis movement disruptsorremoves thethin layerof warmed air that builds up near and about thebody. This air movement leads to loss of totalheat, since heat is transferred from the core ofthe body to rewarm the new colder air, replacingthat blown away. Therefore, wind chill not onlyleads to frostbite locally, but may contribute togeneral hypothermia.

During the antarctic winter of 1941 Sipleand Passe1 developed a formula to determinewind chill from experiments made at LittleAmerica (Siple and Passe1 1945). The formularelates heat loss (H) from an object or person towind speed and to the difference in temperaturebetween the air and the object or person (DT). Itis measured in heat units (calories) per unit areaover time. The skin temperature of most peopleis approximately 33°C (91.4”F). Heat losses forthe human body can then be computed for anyc o m b i n a t i o n o f w i n d a n d t e m p e r a t u r e .Equivalent temperature is based on calm condi-tions and a person walking vigorously at 3 knots(4 mph). Each combination of wind and airtemperature produces a heat loss H. Theequivalent temperature is that temperature thatwould compute the same heat loss at a wind of3 knots. The accompanying chart, figure 18,shows equivalent wind chill temperatures in “Cfor various combinations of winds in knots orkmlhr and temperatures.

Concepts in the following discussion ofwind chill are from an appendix to an article by

Equivalent Wind Chill Temperature

Wllld Speed Coolina Power 01 Wind ExPressed As “Ewwalent Chtll Temperature”

I I Equivalent Chill Temperature I

.: 1 6 14 -21 2i - 34 - 40 -47 - 53 -60 -66 .-73 -79 --66 -92 -99 - 105 - 1 1 2 -118

‘2 ‘,, 1 8 14 -21 27 34 - 40 -47 - 53 -60 -66 -73 7’79 .-.-86 -92 -99 - 1 0 5 -1i2 -.I 18.,_LItlIe Danger Increasing Danger Great Danger- * -

i..,-,” ,j. ‘,_ -/

, .I ‘.,.:/ ,,; :

(Flesh May Freeze- ’ ‘*,,-*p : i(Flesh>May FreGe.

..c

Wblhm 1 Mrwte) ithin- Seconds): ..e-;T.+7..“: kc-., , ~

,i: .:. . , ’ .-. ._ i 0 _ s“‘:,‘; I.,‘.; .__. “,.. .,-

. : ’ ; ;i) ,

Danger of Freeztng Exposed Flesh For Properly Clothed lndlvlduals

Figure 22. Equivalent Wind Chill Temperature Adapted from NWSlNOAA Technical ProceduresBulletin No. 165. Effective Temperature (Wind Chill Index) 1976.

William J. Mills, Jr., M.D., as published in AlaskaMedicine (1973). Dr. Mills is still active in thetreatment of cold injuries in Alaska.

Almost everyone knows that the increasedspeed of wind may cause increased danger ofskin freezing. Many assume that the increase inwind speed causes the ambient air temperaturetofalllower.Thisisnotso.Whatdoesoccurisairmovement, so that warmed air is moved awayfrom the individual exposed to the wind, causingfirst local, then general body cooling. Any result-ant decrease of skin temperature is due to heatloss, insidious or sudden. Local vasoconstric-tion, vascular shunting, and cellular changestake place; eventually ice forms in the tissues,with true tissue freezing or frostbite.

This phenomenon can be readily proved.Place a laboratory recording thermometer with athermistor attached (or any outdoor ther-mometer) out your car window on a calm daywhen the temperature is, say, - 20°C (- 4”F)-just a nice winter day in Anchorage, Alaska. Letit sit for a few minutes until the temperature

reading has stabilized. This temperature, asread, will remain at the ambient air temperaturelevel. Now slowly accelerate your vehicle to80 kmlhr (50 mph); the temperature remainsunchanged at -20°C (-4°F). Now attach thethermistor to your bare hand. Place your un-gloved hand out the same car window in thesame ambient temperature of - 20 “C (- 4 “F).After a few minutes at 0 kmlhr, the skintemperature may be read at approximately 93 “F(normal skin temperature in the nonsmoker).Your skin temperature will drop as heat is lost tothe exterior, sometimes falling as low as 8.5” to80 “F very rapidly. As the car is accelerated andthe warmed air layer is moved away, the ther-mistor records skin heat loss. If you continuedriving, your skin temperature may drop to alevel near 23°F (-5”C), the temperature atwhich freezing of skin may actually occur.

Wind chill may occur not only from naturalwind, but also with air movement generated byautomobile, snowmobile, aircraft, or helicopterrotoblade. These vehicles may predisposepassengers to frostbite or general hypothermia.