us navy course navedtra 14010 - aerographer's mate 1 & c

8/14/2019 US Navy Course NAVEDTRA 14010 - Aerographer's Mate 1 & C

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DISTRIBUTION STATEMENT A: Approved for public release; distribution is unlimited.

NONRESIDENTTRAINING

COURSESeptember 1995

Aerographer's Mate1 & CNAVEDTRA 14010


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DISTRIBUTION STATEMENT A: Approved for public release; distribution is unlimited.

Although the words he, him, andhis are used sparingly in this course toenhance communication, they are notintended to be gender driven or to affront ordiscriminate against anyone.


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PREFACE

By enrolling in this self-study course, you have demonstrated a desire to improve yourself and the Navy.Remember, however, this self-study course is only one part of the total Navy training program. Practicalexperience, schools, selected reading, and your desire to succeed are also necessary to successfully round

out a fully meaningful training program.THE COURSE : This self-study course is organized into subject matter areas, each containing learningobjectives to help you determine what you should learn along with text and illustrations to help youunderstand the information. The subject matter reflects day-to-day requirements and experiences of personnel in the rating or skill area. It also reflects guidance provided by Enlisted Community Managers(ECMs) and other senior personnel, technical references, instructions, etc., and either the occupational ornaval standards, which are listed in the Manual of Navy Enlisted Manpower Personnel Classificationsand Occupational Standards , NAVPERS 18068.

THE QUESTIONS : The questions that appear in this course are designed to help you understand thematerial in the text.

VALUE : In completing this course, you will improve your military and professional knowledge.Importantly, it can also help you study for the Navy-wide advancement in rate examination. If you arestudying and discover a reference in the text to another publication for further information, look it up.

1995 Edition Prepared by AGCM B. J. Bauer and AGC(AW) T. Howlett

Published byNAVAL EDUCATION AND TRAINING

PROFESSIONAL DEVELOPMENTAND TECHNOLOGY CENTER

NAVSUP Logistics Tracking Number0504-LP-026-6930


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Sailors Creed

I am a United States Sailor.

I will support and defend theConstitution of the United States of

America and I will obey the ordersof those appointed over me.

I represent the fighting spirit of theNavy and those who have gonebefore me to defend freedom anddemocracy around the world.

I proudly serve my country s Navycombat team with honor, courageand commitment.

I am committed to excellence andthe fair treatment of all.


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CONTENTS

C H A P T E R Page

1. Convergence, Divergence, and Vorticity . . . . . . . . . . . . . . . . 1-1

2. Forecasting Upper Air Systems . . . . . . . . . . . . . . . . . . . . . . 2-1

3. Forecasting Surface Systems . . . . . . . . . . . . . . . . . . . . . . . . . 3-1

4. Forecasting Weather Elements . . . . . . . . . . . . . . . . . . . . . . . 4-1

5. Forecasting Severe Weather Features . . . . . . . . . . . . . . . . . . 5-1

6. Sea Surface Forecasting . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6-1

7. Meteorological Products and Tactical Decision Aids . . . . . . . 7-1

8. Oceanographic Products and Tactical Decision Aids . . . . . . . 8-1

9. Operational Oceanography . . . . . . . . . . . . . . . . . . . . . . . . . . 9-1

10. Special Observations and Forecasts . . . . . . . . . . . . . . . . . . . 10-1

11. Tropical Forecasting . . . . . . . . . . . . . . . . . . . . . . . . . . . .. .11-1

12. Weather Radar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .. .12-1

13. Meteorological and Oceanographic Briefs . . . . . . . . . . . . . . . 13-1

14. Administration and Training . . . . . . . . . . . . . . . . . . . . . . . . 14-1

A P P E N D I X

I. References Used to Develop the TRAMAN . . . . . . . . . . . . . AI-1

INDEX . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . INDEX-1

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INSTRUCTIONS FOR TAKING THE COURSE

ASSIGNMENTS

The text pages that you are to study are listed atthe beginning of each assignment. Study thesepages carefully before attempting to answer thequestions. Pay close attention to tables andillustrations and read the learning objectives.The learning objectives state what you should beable to do after studying the material. Answeringthe questions correctly helps you accomplish theobjectives.

SELECTING YOUR ANSWERS

Read each question carefully, then select theBEST answer. You may refer freely to the text.The answers must be the result of your ownwork and decisions. You are prohibited fromreferring to or copying the answers of others andfrom giving answers to anyone else taking thecourse.

SUBMITTING YOUR ASSIGNMENTS

To have your assignments graded, you must beenrolled in the course with the NonresidentTraining Course Administration Branch at theNaval Education and Training ProfessionalDevelopment and Technology Center(NETPDTC). Following enrollment, there aretwo ways of having your assignments graded:(1) use the Internet to submit your assignmentsas you complete them, or (2) send all theassignments at one time by mail to NETPDTC.

Grading on the Internet: Advantages toInternet grading are:

you may submit your answers as soon asyou complete an assignment, and

you get your results faster; usually by thenext working day (approximately 24 hours).

In addition to receiving grade results for eachassignment, you will receive course completionconfirmation once you have completed all the

assignments. To submit your assignmentanswers via the Internet, go to:

http://courses.cnet.navy.mil

Grading by Mail: When you submit answersheets by mail, send all of your assignments atone time. Do NOT submit individual answersheets for grading. Mail all of your assignmentsin an envelope, which you either provideyourself or obtain from your nearest EducationalServices Officer (ESO). Submit answer sheetsto:

COMMANDING OFFICERNETPDTC N3316490 SAUFLEY FIELD ROADPENSACOLA FL 32559-5000

Answer Sheets: All courses include onescannable answer sheet for each assignment.These answer sheets are preprinted with yourSSN, name, assignment number, and coursenumber. Explanations for completing the answersheets are on the answer sheet.

Do not use answer sheet reproductions: Useonly the original answer sheets that weprovidereproductions will not work with ourscanning equipment and cannot be processed.

Follow the instructions for marking youranswers on the answer sheet. Be sure that blocks1, 2, and 3 are filled in correctly. Thisinformation is necessary for your course to beproperly processed and for you to receive creditfor your work.

COMPLETION TIME

Courses must be completed within 12 monthsfrom the date of enrollment. This includes timerequired to resubmit failed assignments.


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PASS/FAIL ASSIGNMENT PROCEDURES

If your overall course score is 3.2 or higher, youwill pass the course and will not be required toresubmit assignments. Once your assignmentshave been graded you will receive coursecompletion confirmation.

If you receive less than a 3.2 on any assignmentand your overall course score is below 3.2, youwill be given the opportunity to resubmit failedassignments. You may resubmit failedassignments only once. Internet students willreceive notification when they have failed anassignment--they may then resubmit failedassignments on the web site. Internet studentsmay view and print results for failedassignments from the web site. Students who

submit by mail will receive a failing result letterand a new answer sheet for resubmission of eachfailed assignment.

COMPLETION CONFIRMATION

After successfully completing this course, youwill receive a letter of completion.

ERRATA

Errata are used to correct minor errors or delete

obsolete information in a course. Errata mayalso be used to provide instructions to thestudent. If a course has an errata, it will beincluded as the first page(s) after the front cover.Errata for all courses can be accessed andviewed/downloaded at:

ht tp : / /www.advancement .cnet .navy.mil

STUDENT FEEDBACK QUESTIONS

We value your suggestions, questions, and

criticisms on our courses. If you would like tocommunicate with us regarding this course, weencourage you, if possible, to use e-mail. If youwrite or fax, please use a copy of the StudentComment form that follows this page.

For subject matter questions:

E-mail: [email protected]: Comm: (850) 452-1001, Ext. 1713

DSN: 922-1001, Ext. 1713FAX: (850) 452-1370(Do not fax answer sheets.)

Address: COMMANDING OFFICERNETPDTC (CODE 315)6490 SAUFLEY FIELD ROADPENSACOLA FL 32509-5237

For enrollment, shipping, grading, orcompletion letter questions

E-mail: [email protected]: Toll Free: 877-264-8583

Comm: (850) 452-1511/1181/1859

DSN: 922-1511/1181/1859FAX: (850) 452-1370(Do not fax answer sheets.)

Address: COMMANDING OFFICERNETPDTC (CODE N331)6490 SAUFLEY FIELD ROADPENSACOLA FL 32559-5000

NAVAL RESERVE RETIREMENT CREDIT

If you are a member of the Naval Reserve, youwill receive retirement points if you are

authorized to receive them under currentdirectives governing retirement of NavalReserve personnel. For Naval Reserveretirement, this course is evaluated at 8 points.(Refer to Administrative Procedures for Naval

Reservists on Inactive Duty, BUPERSINST1001.39, for more information about retirementpoints.)

COURSE OBJECTIVES

The objective of this course is to provide

Aerographer's Mates with occupational informa-tion on the following areas: convergence,divergence, and vorticity; the forecasting of upper air systems; the forecasting of surfacesystems; the forecasting of weather elements;the forecasting of severe weather features; seasurface forecasting; meteorological products andtactical decision aids; oceanographic productsand tactical decision aids; operational oceano-


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graphy; tropical forecasting; weather radar;meteorological and oceanographic briefs; andadministra-tion and training.


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Student Comments

Course Title: Aerographer's Mate 1 & C

NAVEDTRA: 14010 Date :

We need some information about you :

Rate/Rank and Name: SSN: Command/Unit

Street Address: City: State/FPO: Zip

Your comments, suggestions, etc .:

Privacy Act Statement: Under authority of Title 5, USC 301, information regarding your military status isrequested in processing your comments and in preparing a reply. This information will not be divulged withoutwritten authorization to anyone other than those within DOD for official use in determining performance.

NETPDTC 1550/41 (Rev 4-00)


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CHAPTER 1

CONVERGENCE, DIVERGENCE, AND VORTICITY

In your reading of the AG2 manual, volume 1, youbecame famil iar with the terms c o n v e r g e n c e ,divergence, and vorticity when used in relation tosurface lows and highs. You were also presented witha basic understanding of the principles involved In thissection, we will cover the terms, the motions involvedin upper air features and surface features, and therelationship of these processes to other meteorologicalapplications,

We will first discuss convergence and divergence,followed by a discussion of vorticity.

N O T E

T h e Wo r l d M e t e o r o l o g i c a lOrgan ization adopted hectopascals( h P a ) a s i t s s t a n d a r d u n i t o f measurement for pressure. Becausethe units of hectopascals a nd m illibarsare interchangeable (1 hPa = 1 mb),hectopascals have been substituted formillibars in this TRAMAN.

CONVERGE NCE ANDD I V E R G E N C E

LEARNING OBJECTIVES: Define the termsconvergence and divergence . R e c o g n i z edirectional and velocity wind shear rules.Recognize areas of mass divergence and massconvergence on surface pressure charts.Identify the isopycnic level. Retail the effectsthat convergence and divergence have onsurface pressure systems and features aloft.Identify rules associated with divergence andconvergence.

As mentioned in the AG2 manual, volume 1, unit 8,convergence is the accumulation of air in a region orlayer of the atmosphere, while divergence is thedepletion of air in a region or layer. The layer of maximum convergence and divergence occurs between

the 300- and 200-hPa levels. Coincidentally, this is alsothe layer of maximum winds in the a tmosphere; where

jet st ream cores a re u sually found. Th ese high-speedwinds are direct ly re la ted to convergence anddivergence. The combined effects of wind direction andwind speed (velocity) is what produces convergent anddivergent airflow.

C O N V E R G E N C E A N D D I V E R G E N C E( S I M P L E M O T I O N S )

Simply stated, convergence is defined as theincrease of mass within a given layer of the atmosphere,while divergence is the decrease of mass within a given

layer of the atmosphere.

Convergence

For convergence to take place, the winds mu st r esultin a net inflow of, air into that layer. We generallyassociate this type of convergence with low-pressureareas, where convergence of winds toward the center of the low results in an increase of mass into the low andan upward m otion. In meteorology, we distinguishbetween two types of convergence as either horizontalor vertical convergence, depending upon the axis of the

flow.

Dive rgence

Winds in t his situ at ion pr oduce a n et flow of airoutward from the layer. We associate this type of divergence with high-pressure cells, where the flow of air is directed outward from the center, causing adownward motion. Divergence, too, is classified aseither horizontal or vertical.

D I R E C T I O N A L W I N D S H E A R

The simplest forms of convergence and divergenceare the types that result from wind direction alone. Twoflows of air need not be moving in opposite directionsto induce divergence, nor moving toward the same pointto induce convergence, but maybe at any angle to eachother to create a net inflow of air for convergence or anet outflow for divergence.

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W I N D S P E E D ( VE L O C I T Y ) S H E A R

Convergence is occurring when wind speeds aredecreasing downstream; that is, mass is accumulatingupstream. Conversely, divergence is occurring whenwind speeds are increasing downstream; that is, mass isbeing depleted upstream.

D I R E C T I O N A L A N D S P E E D WI N D S H E A R

Wind speed in relation to the wind direction is alsoa valuable indicator. For example, on a streamlineana lysis chart we can ana lyze both wind direction a ndwind speed, variations in wind speed along thestreamlines, or the convergence or divergence of thestreamlines.

The following are some of the combinations or

variations of wind speed and direction:

. In a field of para llel streamlines (wind flow), if the wind speed is decreasing downstream (producing anet inflow of air for th e layer), convergence is tak ingplace. If the flow is increasing downstream (a netoutflow of air from the layer), divergence is occurring.

. In an area of uniform wind speed along thestreamlines, if the streamlines diverge (fan out),divergence is occurring; if the streamlines converge(come together), convergence is taking place.

. Normally, the convergence and divergencecomponent s ar e combined. The fact th at str eamlinesconverge or diverge does not necessarily indicateconvergence or divergence. We must also consider thewind speeds whether they are increasing ordecreasing downstream in relation to whether thestreamlines are spreading out or coming together.

. If, when looking downstream on the streamlines,the wind speed increases and the streamlines diverge,divergence is ta king place. On th e other h and, if thewind speed decreases downstream and the streamlinescome together, convergence is taking place.

There are other situations where it is more difficultto determine whether divergence or convergence isoccurring, such as when the wind speed decreasesdownst ream and the wind flow diverges, as well as whenwind speed increases downstream and the wind flowconverges. A special evaluation then must be made todetermine the net inflow or outflow.

D I V E R G E N C E A N D C O N V E R G E N C E(COMPLEX MOTIONS)

In this section we will be discussing high-levelconvergence and divergence in relation to downstreamcontour patterns and the associated advection patterns.Low tropospheric advection (and also stratosphericadvection) certainly play a large role in pressure changemechanisms.

Since the term divergence is meant to denotedepletion of mass, while convergence is meant to denoteaccumulation of mass, the forecaster is concerned withthe mass divergence or mass convergence in estimatingpressure or height changes. Mass divergence in theentire column of air produces pressure or height falls,while mass convergence in the entire column of airproduces pressure or height rises at the base of thecolumn.

Mass Dive rgence and Mass Conve rgence

Mass divergence and mass convergence involve thedensity field as well as the velocity field. However, themass divergence and mass convergence of theatmosphere are believed to be largely stratified into twolayers as follows:

l Below about 600 hPa, velocity divergence andconvergence occur chiefly in the friction layer, which isabout one-eighth of the weight of the 1,000-to 600-hPaadvect ion s t ra tum, and may be disregarded incomparison with density transport in estimating thecontribut ion t o the pressu re chan ge by th e advectionstratum.

. A b o v e 6 0 0 h P a , m a s s d i v e rg e n c e a n dconvergence largely result from horizontal divergenceand convergence of velocity. However, on occasion,stratospheric advection of density may be a modifyingfactor.

The stratum below the 400-hPa level may beregarded as the ADVECTION stratum, while thestratum above the 400-mb level maybe regarded as thesignificant horizontal divergence or convergence

stra tum . Also, the advection st rat um m aybe thought of as the zone in which compensation of the dynamiceffects of the upper stratum occurs.

The Isopycnic Level

At about 8km (26,000 ft) the density is nearlyconstant. This level, which is near the 350-hPa pressuresurface, is called the isopycnic level. This level is th e

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location of constant density, with mass variations aboveand below.

Since the density at 200 hPa is only four-seventhsthe density at the isopycnic level, the height change at200 hPa would have to be twice th at at th e isopycniclevel (350 hPa) for the same pressure/height change tooccur. Thus, height changes in the lower stratospheretend to be a ma ximum even th ough pressure changes area maximum at the isopycnic level.

Pressure changes occur at the isopycnic level, andin order to maintain constant density a correspondingtemperat ure chan ge must also occur. Since the densityis nearly constan t a t t his level, the r equired temperaturevariations must result from vertical motions. When thepressures are rising at this level, the temperature mustalso rise to keep the density constant. A temperature risecan be produced by descending motion. Similarlyfa l l ing pressures a t this level require fa l l ingtemperatures to keep the density constant. Fallingtemperatures in the absence of advection can be

produced by ascent through this level.

Thus, rising heights at the isopycnic level areassociated with subsidence, and falling heights at theisopycnic level are associated with convection.

T h e 3 5 0-h P a t o 2 0 0-h P a S t r a t u m

Subsidence at 350 hPa can result from horizontalconvergence above this level, while convection herewould result from horizontal divergence above thislevel.

Since rising heights in the upper troposphere resultin a rising of the tropopause and the lower stratosphere,the maximum horizontal convergence must occurbetween the isopycnic level (350 hPa) and the averagelevel of the tropopause (about 250 hPa). This is due tothe reversal of the vertical motion between thetr opopau se an d th e isopycnic level. Thu s, th e level of maximum horizontal velocity convergence must bebetween 300 hPa and 200 hPa and is the primarymechanism for pr essure or height rises in th e upper a ir.Similarly, upper height falls are produced by horizontalvelocity divergence with a maximum at the same level.The maximum divergence occurs near or slightly abovethe tropopause and closer to 200 hPa than to 300 hPa.Therefore, it is more realistic to define a layer of maximum divergence and convergence as occurringbetween the 300- and 200-hPa pressure surfaces. The300- to 200-hPa stratum is also the layer in which thecore of the jet stream is usually located. It is also at thislevel that the cumulative effects of the mean temperature

field of the troposphere produce the sharpest horizontalcontrasts in the wind field.

The level best sui ted for determinat ion of convergence and divergence is the 300-hPa level.Because of the sparsity of reports at the 300-hPalevel, it is frequently advant ageous t o determ ine th epresence of convergence and divergence at the 500-hPalevel.

Dive rgence /Conve rgence and Sur faceP r e s s u r e S ys t e m s

The usua l d i s t r ibu t ion o f d ivergence andconvergence relative to moving pressure systems is asfollows:

l In advance of the low, convergence occurs at lowlevels and divergence occurs aloft, with the level of nondivergence at about 600 hPa.

l In the rear of the low, there is usual ly

convergence aloft and divergence near the surface.The low-level convergence ahead of the low occurs

usually in the stratum of strongest warm advection, andthe low-level divergence in the rear of the low occurs inthe stratum of strongest cold advection. The low-leveldivergence occurs primarily in the friction layer(approximately 3,000 ft) and is thought to be of minorimportance in the modification of thickness advectioncompared with heating and cooling from the underlyingsurfaces.

Dive rgence /Conve rgen ce Fea t u r e s Alo ft

In a dvance of the low, the air r ises in response to thelow-level convergence, with the maximum ascendingmotion at the level of nondivergence eventuallybecoming zero at the level of maximum horizontaldivergence (approximately 300 hPa). Above this level,descending motion is occurring. In the rear of the low,th e reverse is true; tha t is, descending motion in t hesurface stratum and ascending motion in the uppertr oposphere above th e level of maximu m horizont alconvergence. In deepening systems, the convergence

aloft to the rear of the low is small or may even benegative (divergence). In fil l ing systems, thedivergence aloft in advance of the low is small or evennegative (convergence).

Thus, in the development and movement of surfacehighs a nd su rface lows, two vert ical circulat ions ar einvolved, one below and one above the 300-hPa level.The lower vertical circulation is upward in the cyclone

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Figure 1-1.-Genera l ized ve r t i ca l c ir cu la t ion overd eve lop ing h ighs a nd lows .

and downward in the anticyclone. The upper verticalc i rcu la t ion invo lves downward mot ion in thestratosphere of the developing cyclone and upward

motion in the upper troposphere and lower stratosphereof the developing anticyclone. See figure 1-1.

Divergence and upper-height falls are associatedwith high-speed winds approaching weak contourgradients which are cyclonically curved. Figure 1-2illustrates contour patterns associated with height falls,

Convergence and upper-height rises are associatedwith the following:

. Low-speed winds approaching straight orcyclonically curved strong contour gradients. Seefigur e 1-3 , view (A).

. High-speed winds approaching anticyclonicallycurved weak contour gradients.(B).

Se e figur e 1-3, view

Figu re 1-2.-Dive rgence I llus t r a t ed .

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Figure 1-3.-Convergence I l lus t r a t ed .


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Note that the associated height rises or falls occurdownstream and to the left of the flow, as illustrated infigure 1-3.

Divergence Iden t i f ica t ion (Downs t r ea mSt ra igh t l i ne F low)

The technique for determining the areas of divergence consists in noting those areas where windsof high speed are approaching weaker downstreamgradients that are straight. When inertia carries ahigh-speed parcel of air into a region of weak gradient,it possesses a Coriolis force too large to be balanced byth e weaker gra dient force, It is thu s deflected t o theright. This r esults in a deficit of ma ss to th e left. Theparcels that are deflected to the right must penetratehigher pressure/heights and are thus slowed down untilthey are in balance with th e weaker gradient. Then theycan be steered along the existing isobaric or contourchannels.

Dive rgence Iden t i f i ca t ion (Weak Downs t r eamCyclonica l ly Curved Flow)

If the weak downstream gradients are cyclonicallycurved, the divergence resulting from the influx of high-speed wind is even more marked due to theadditional effect of centrifugal forces.

Dive rgence Iden t i f ica t ion (Downs t r ea mAnt icyclonica l ly Curved Flow)

The effect of centrifugal forces on anticyclonically

curving high-speed parcels is of extreme importance inproducing overshooting of high-speed air from sharplycurved ridges into adjacent troughs, causing pressurerises in the west side of the troughs.

Dive rgence Iden t i fi ca t ion (S t rong Winds )

I f h igh-speed parce l s approach d iverg ingcyclonically curved cont ours, lar ge cont our falls willoccur downstream to the left of the high-speed winds.Eventually a strong pressure gradient is produceddownstream, to the right of the high-speed winds,

chiefly as a result of pressur e falls to the left of th edirection of high-speed winds in the cyclonically curvedcontours with weak pressure gradient. Usually thedeflection of air toward higher pressure is so slight thatit is hardly observable in individual wind observations.However, when the pressure field is very weak to thetight of the incoming h igh-speed str eam, n oticeableangles between the wind and contours may be observed,

especially at lower levels, due to tr an sport of moment umdownward as a result of subsidence, where the gradientsare even weaker . This occurs sometimes t o such anextent that the wind flow is considerably more curvedant icyclonically tha n t he cont our s. In r are cases th isresults in anticyclonic circulation centers out of phasewith the high-pressure center. This is a transitorycondition necessitating a migration of the pressurecenter toward the circulation center. In cases where thehigh-pressure center and anticyclonic wind flow centerare out of phase, the pressure center will migrate towardthe circulation center (which is usually a center of massconvergence).

It is more normal, however, for the wind componenttoward high pressure to be very slight, and unless thewinds and contours are drawn with great precision, thedeviation goes unnoticed.

Overshoo t ing

High-speed winds approaching sharply curvedridges result in large height rises downstream from theridge due to overshooting of the high-speed air. It isknown from the gradient wind equation that for a givenpressure gradient there is a limiting curvature to thetrajectory of a parcel of air moving at a given speed.Frequently on upper air charts, sharply curvedstationary ridges are observed with winds of high speedapproaching the r idge. The existence of a sharplycurved extensive ridge usually means a well-developedtrough downstream, and frequently a cold or cutoff low

exists in this trough. The high-speed winds approachingthe ridge, due to centrifugal forces, are unable to makethe sharp turn necessary to follow the contours. Thesewinds overshoot the ridge anticyclonically, but with lesscurvatur e tha n t he contours, resulting in their plungingacross contours toward lower pressure/heightsdownstrea m from the ridge. This may r esult in an yoneof a number of consequences for the downstreamtr ough, depending on t he init ial configur at ion of theridge and t rough, but a ll of these consequen ces are basedon the convergence of mass into the trough as a resultof overshooting of winds from the ridge.

Four effects of overshooting areas follows:

1. Filling of the downstream trough. This happensif the contour gradient is strong on the east side of thetrough; tha t is, a blocking ridge to the eas t of the tr ough.

2. Acceleration of the cutoff low from of itsstationary position. This usually occurs in all cases.

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3. Radical reorientation of the trough. This usuallyhappens where the trough is initially NE-SW, resultingin a N-S and in some cases a N W-SE orientat ion a ftersufficient tim e (36 hours).

4. This situation may actually cut off a low in thelower area of the trough. This usually happens when thehigh-speed winds approach ing the r idge a resouthwesterly and approach the ridge at a comparativelyhigh latitude relative to the trough. This frequently

reorients the trough line towards a more NE-SWdirection. Usually, the reorientation of the troughoccurs simultan eously with 1 and 2.

Sharp ly Curved R idges

Closely related to the previously mentionedsituation are cases of sharply curved ridges where thegradient in the sharply curved portion (usually thenorthern port ions of a north-south r idge) hasmomentarily built up to a strength that is incompatiblewith the anticyclonic curvature. Such ridges oftencollapse with great rapidity prior to the development of such excessive gradients, causing rapid filling of theadjacent downstream trough, and large upper contourfa l ls . The grad ien t wind re la t ion impl ies tha tsubsequent trajectories of the high-speed parcelsgenerated in the strong ridge line gradient must be lessanticyclonically curved than the contours in the ridge.

It can also be shown from the gradient windequation that the anticyclonic curvature increases as thedifference between the actual wind and the geostrophicwind increases, until the actual wind is twice the

geostrophic wind, when the trajectory curvature is at amaximum . This fact can be used in determining th etrajectory of high-speed parcels approaching sharplycurved stationary ridges or sharply curved stationaryridges with strong gradients. By measuring thegeostrophic wind in the ridge, the maximum trajectorycurvature can be obtained from the gradient wind scale.This trajectory curve is the one that an air parcel at th eorigin point of the scale will follow until it intersects thecorrection curve from the geostrophic speed to thedisplacement curve of twice the geostrophic speed.

Actua l Wind Speeds

If actual wind speed observations are available forparcels approaching the ridge, comparison can be madewith the geostrophic winds (pressure gradient) in theridge. If the actual speeds are more than twice themeasu red geostrophic wind in t he r idge, the an ticycloniccurvature of these high-speed parcels will be less than

the maximum trajectory curvature obtained from thegradient wind scale, and even greater overshooting of these high-speed parcels will occur across lowercontours . Convergence in the west s ide of thedownstream trough results in lifting of the tropopausewith dynamic cooling and upper-level contour rises.

Subgrad ien t Winds

Low-speed winds approaching an area of strongergradient become subject to an unbalanced gradient forcetoward the left due to the weaker Coriolis force. Thesesubgradient winds are deflected toward lower pressure,crossing contours a nd producing cont our r ises in t hearea of cross-contour flow. This cross-contour flowaccelerates the air until it is moving fast enough to bebalanced by the st ronger pr essure gradient. Due t o theacceleration of the slower oncoming parcels of air, thecontour rises propagate much faster than might beexpected on the basis of the slow speed of the air as it

initially enters the stronger pr essure gradient.

The following two rules summarize the discussionof subgradient winds:

. High-speed winds approaching low-speed windswith weak cyclonically curved contour gradients areindicative of divergence and upper-height fallsdownstream to the left of the current.

. L o w - s p e e d w i n d s a p p r o a c h i n g s t r o n g ,cyclonically curved contour gradients or high-speed

winds approaching low-speed winds with weak anticyclonically curved contour gradients are indicativeof convergence and upper height rises downstream andto the left and right of the current, respectively.

I M P O RTA N C E O F C O N VE R G E N C E AN DD I V E R G E N C E

Convergence and divergence have a pronouncedeffect upon the weather occurring in the atmosphere.Vertical motion, either upward or downward, isrecognized as an important parameter in theatmosphere. For instance, extensive regions of precipitation associated with extratropical cyclones areregions of large-scale upward motion. Similarly, thenearly cloud-free regions in large anticyclones areregions in which air is subsiding. Vertical motions alsoaffect temperature, humidity, and other meteorologicalelements.

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Chan ges in S t ab i l i t y

When convergence or divergence occurs, whetheron a large or small scale, it may have a very pronouncedeffect on th e sta bility of the a ir. For example, whenconvection is in duced by conver gence, air is forced t orise without the addition of heat. If this air isunsaturated, it cools first at the dry adiabatic rate; or if satura ted at t he moist ra te. The end result is that the airis cooled, which will increase the instability y of that aircolumn due to a net release of heat. Clouds and weatheroften result from this process.

Conversely, if air subsides, and this process isproduced by convergence or divergence, the sin king a irwill heat at the dry adiabatic lapse rate due tocompression. The warming at the top of an air columnwill increase the stability of that air column by reducingthe lapse r ate. Such war ming often dissipates existingclouds or prevents the formation of new clouds. If sufficient warming due to the downward motion takesplace, a subsidence inversion is produced.

Effec t on Wea th e r

The most important application of vertical motionis the prediction of rainfall probability and rainfallamount. In addition, vertical motion affects practicallyall meteorological properties, such as temperature,hum idity, wind distribution, and par ticularly sta bility.In the following section the distribution of large-scaleand small-scale vertical motions are considered.

Since cold air has a tendency to sink, subsidence islikely to be found to th e west of upper tr opospherictroughs, and rising air t o the east of the t roughs. Thus,there is a good relation between upper air meridionalflow and vertical flow.

In the neighborhood of a straight NorthernHemisphere jet stream, convergence is found to thenorth of the stream behind centers of maximum speedas well as to the south and ahead of such centers.Divergence exists in the other two quadrants. Below theregions of divergence the air rises; below those of convergence there is subsidence.

These general rules of thumb are not perfect, andonly yield a very crude idea about dist ribution of verticalmotion in the horizontal. Particularly over land insummer, there exists little relation between large-scraleweather patterns and vertical motion. Rather, verticalmotion is influenced by local features and shows strongdiurn al var iations. Large-scale vertical motion is of small magnitude at the ground (zero if the ground is flat).

Above ground level, it increases in magnitude to at least500 hPa and decreases in the neighborhood of thetropopause. There have been several studies of therelation between frontal precipitation and large-scalevertical velocities, computed by various techniques. Inall cases, the probability of precipitation is considerablyhigher in the 6 hours following an updra ft tha n followingsubs idence . C lea r sk ies a re mos t l ike ly wi thdowndra fts. On t he other ha nd, it is not obvious that

large-scale vertical motion is relat ed to showers a ndthunderstorms caused during the daytime by heating.However, squall lines, which are formed along lines of horizontal convergence, show that large-scale verticalmotion may also play an important part in convectiveprecipitation.

Ver t i ca l Ve loci ty Cha r t s

Vertical velocity charts are currently being

transmitted over the facsimile network and arecomput ed by numer ical weath er pr ediction meth ods.The charts have plus signs indicating upward motionand minus signs indicating downward motion. Thefigures indicate vertical velocity in centimeters persecond (cm/sec). With the larger values of upwardmotions (plus values) the likelihood of clouds andprecipitation increases. However, an evaluation of themoisture and vertical velocity should be made to getoptimum results. Obviously, upward motion in dry airis not as likely to produce precipitation as upward

motion in moist air.

Studies have shown that surface cyclones andanticyclones are not independent of developments in theupper at mosphere, rather, t hey work in tandem with oneanother. The relationship of the cyclone to thelarge-scale flow patterns aloft must therefore be a partof the daily forecast routine.

Many forecasters have a tendency to shy away fromthe subject of vorticity, as they consider it too complexa subject to be mastered. By not considering vorticity

and its effects, the forecaster is neglecting an importantforecasting tool. The principles of vorticity are no morecomplicated than most of the principles of physics, andcan be understood just as readily. In the followingsection we will discuss th e definition of vorticity, itsevaluation, and its relationships to other meteorologicalparameters.

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VORTICITY

LEARN ING OBJECTIVES : Recognize t he twocomponents of relative vorticity. Define theterm absolute vorticity. Determine vorticityimpacts on weather processes.

Vort icity measu res t he r otation of very small airparcels. A parcel has vorticity when it spins on its axisas it moves along its path. A parcel that does not spinon its axis is said to have zero vorticity. The axis of spinning or rotation can extend in any direction, but forour purposes, we are mainly concerned with therotational motion about an axis that is perpendicular tothe surface of the Earth. For example, we could drop achip of wood into a creek and watch its progress. Thechip will move downstream with the flow of water, butit ma y or m ay not spin as it moves downstr eam. If itdoes spin, the chip has vorticity. When we try to isolatethe cause of the spin, we find that two properties of theflow of wat er cau se t he chip t o spin: (1) If th e flow of water is moving faster on one side of the chip than theother, this is shear of the current; (2) if the creek bedcurves, the path has curvature. Vorticity always appliesto extremely small air parcels; thus, a point on one of our upper air charts may represent such a parcel. Wecan examine this point and say that the parcel dots ordoes not have vorticity. However, for this discussion,larger parcels will have to be used to more easilyvisualize the effects. Actually, a parcel in theatmosphere has three rotational motions at the sametime: (1) rotation of the parcel about its own axis(shear), (2) rotation of the parcel about the axis of apressure system (curvature), an d (3) rotat ion of theparcel due to the atmospheric rotation. The sum of thefirst two components is known as relative vorticity, andthe sum total of all three is known as absolute vorticity.

R E L AT I V E V O RT I C I T Y

Relative vorticity is the sum of the rotation of theparcel about the axis of the pressure system (curvature)and the rotation of the parcel about its own axis (shear).

Figure 1-4.-I l lus t ra t ion of vor t ic i ty due to th e shea r effec t .

The vorticity of a horizonta l curren t can be broken downinto two components, one due to curvature of thestreamlines and the other due to shear in the current.

S h e a r

First, let us examine the shear effect by looking atsmall air parcels in an upper air pat tern of straightcont ours. Here th e wind shear r esults in each of th e

three parcels having different rotations (fig. 1-4).Refer to figure 1-4. Parcel No. 1 has stronger wind

speeds to its right. As the parcel moves along, it will berotated in a counterclockwise direction. Parcel No. 2has the str onger wind speeds to its left; therefore, it willrotate in a clockwise direction as it moves along. ParcelNo. 3 has equal wind speeds to the right and left. It willmove, but it will not rotate. It is said to have zerovorticity.

Therefore, to briefly review the effect of shear-aparcel of the atmosphere has vorticity (rotation) when

th e wind speed is stronger on one side of the par cel tha non the oth er.

Now lets define positive and negative vorticity interms of clockwise and counterclockwise rotation of aparcel. The vorticity is positive when the parcel has acounterclockwise rotat ion (cyclonic , NorthernHemisphere) and the vorticity is negative when theparcel has clockwise rotation (anticyclonic, NorthernHemisphere).

Thus, in figure 1-4, parcel No. 1 has positivevorticity, and parcel No. 2 has negative vorticity.

C u r v a t u r e

Vort icity can also result du e to cur vatu re of theairflow or path. In the case of the wood chip flowingwith the stream, the chip will spin or rotate as it movesalong if the creek curves.

To demonstrate the effect of curvature, let usconsider a patt ern of contours having curvat ure but noshear (fig. 1-5).

Figure 1-5.- I l lus t ra t ion of vor t ic i ty due to curvature ef

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Place a small parcel at the trough and ridge lines andobserve the way the flow will spin the parcel, causingvorticity. The diameter of the parcel will be rotated fromthe solid line to the dotted position (due to the northerlyand southerly components of the flow on either side of the trough and ridge lines).

Note that we have counterclockwise rotation at thetrough (positive vorticity), and at the ridge line we haveclockwise rotation (negative vorticity). At the pointwhere t here is n o curvat ure (inflection point), there is noturning of the parcel, hence no vorticity. This isdemonstrated at point Pin figure 1-5.

Combined Effec t s

To find the relative vorticity of a given parcel, wemust consider both the shear and curvature effects. It isquite possible to have two effects counteract each other;that is, where shear indicates positive vorticity butcurvature indicates negative vorticity, or vice versa (fig.1-6).

To find th e net r esult of th e two effects we wouldmeasur e the value of each and a dd them algebraically.The measurement of vorticity will be discussed in thenext section.

It must be emphasized here that relative vorticity isobserved instantaneously. Relative vorticity in theatmosphere is defined as the instantaneous rotation of very small part icles. The r ota tion results from windshear and curvat ure. We refer t o this vorticity as beingrelative, because all the motion illustrated was relativeto the surface of the Earth.

A B S O L U T E V O RT I C I T Y

When the relative vorticity of a parcel of air isobserved by a person completely removed from theEarth, he or she observes an additional component of vorticity created by the rotation of the Earth. Thus, this

Figur e 1-6-I l lus t r a t ion o f shea r e ffec t oppos ing the cu rva tu r eeffect in producing vorticity. (A) Negative shear and positive

curvature ; (B) pos i t ive shear and negat ive curvature .

Figu re 1-7.-Con tour -i so tach pa t t e r n fo r shea r ana lysi s .

person sees the total or absolute vorticity of the sameparcel of air.

The total vorticity, tha t is, relat ive vorticity plus th atdue to the Ea rt hs rotat ion, is known as the a bsolutevorticity. As was stated before, for practical use inmeteorology, only the vorticity about an axisperpendicular to the surface of the Earth is considered.In t his case, the vort icity due to the E ar th s rotat ionbecomes equal to the Coriolis parameter. This isexpressed as 2oI sin , where w is the angular velocityof the Earth and is the latitude. Therefore, theabsolute vorticity is equal to the Coriolis parameter plusthe relative vorticity. Writing this in equation formgives: (Za = a bsolute vort icity)

Z a = 2 c o s i n 0 + Z r

E VAL U AT I O N O F V O RT I C I T Y

In addition to locating the areas of convergence anddivergence, we must also consider the effects of horizontal wind shear as it affects the relative vorticity,

and hence the movement of the long waves anddeepening or falling associated with this movement.

The two terms c u r v a t u r e and shear, w h i c hdetermine the relative vorticity, may vary inversely toeach other. Therefore, it is necessary to evaluate bothof them. Figures 1-7 through 1-10 illustrate some of thepossible combinat ions of curvatu re and shear . Solid

Figur e 1-8.-Con tour -i so tach pa t t e r n fo r shea r ana lysi s .

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Figure 1-9.-Contour -iso tachch pa t tern for shea r an alys is .

lines are streamlines or contours; dashed lines areisotachs.

Figure 1-7 represents a symmetrical sinusoidalstreamline pattern with isotachs parallel to contours.Therefore, there is no gradient of shear along thecontours. In region I, the curvature becomes moreanticyclonic downstream, reaching a maximum at theaxis of the downstream ridge; that is, relative vorticitydecreases from the trough to a minimum at thedownstream ridge. The region from the trough to thedownstream ridge axis is favorable for deepening.

The reverse is tr ue west of th e tr ough, region II.This region is unfavorable for deepening.

In figure 1-8 there is no curvature of streamlines;therefore, the shear alone determines the relativevort icity. The shear downstr eam in r egions I a nd IVbecomes less cyclonic; in regions II and III, it becomesmore cyclonic. Regions I and IV are therefore favorablefor deepening downstream.

In region I of f igure 1-9 b oth cyclonic shear andcurvature decrease downstream and this region is highlyfavorable for deepening. In region III both cyclonicshear and curvature increase downstream and thisregion is u nfavora ble for deepening. In r egion II t hecyclonic curvature decreases downstream, but thecyclonic shear increases. This situation is indeterminatewithout calculation unless one term predominates. If the curvature gradient is large and the shear gradientsmall, the region is likely to be favorable for deepening.

Figur e 1-10.-Con tour -i so tach pa t t e r n fo r shea r ana lysi s .

In region IV, the cyclonic curvature increasesdownstream, but the cyclonic shear decreases, so thatthis region is also indeterminate unless one of the twoterms predominates.

In region I of figure 1-10 the cyclonic sheardecreases downstream and the cyclonic curvatureincreases. The region is indeterminate; however, if theshear gradient is larger than the curvature gradient,deepening is favored. Region II h as increasing cyclonicshear and curva ture downs t ream and i s qu i teunfavorable. In region III, the shear becomes morecyclonic downstream and the curvature becomes lesscyclonic. This region is also indeterminate unless thecurvature term predominates. In region IV, the shearand curvature become less cyclonic downstream and theregion is favorable for deepening.

R E L AT I O N O F V O RT I C I T Y TO WE AT H E RP R O C E S S E S

Vorticity not only affects the formation of cyclonesand ant icyclones, but it also has a direct bear ing oncloudiness, precipitation, pressure, and height changes.Vorticity is used primarily in forecasting cloudiness andprecipitation over an extensive area. One rule states thatwhen r elative vorticity decreases downstrea m in t heupper troposphere, convergence is taking place in thelower levels . When convergence takes place,cloudiness and possibly precipitation will prevail if sufficient moisture is present.

One rule using vorticity in relation to cyclonedevelopment stems from the observation that whencyclone development occurs, the location, almostwithout exception, is in advance of art upper trough.Thus, when an upper level trough with positive vorticityadvection in advance of it overtakes a frontal system inthe lower troposphere, there is a distinct possibility of cyclone development at the surface. This is usuallyaccompanied by deepening of the sur face system . Also,the development of cyclones at sea level takes placewhen and where an area of positive vorticity advectionsitua ted in t he upper troposphere overlies a slow movingor quasi-stationary front at the surface.

The relationship between convergence anddivergence can best be illustr ated by th e term shear. If we consider a flow where the cyclonic shear isdecreasing downstream (stronger wind to the right thanto the left of the current), more air is being removed fromthe a rea t han is being fed into it, hence a net depletionof mass aloft, or divergence. Divergence aloft isassociated with surface pressure falls, and since this is

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the situation, the relative vorticity is decreasingdownstream. We may state that surface pressure fallswhere relative vorticity decreases downstream in theupper troposphere, or where advection of more cyclonicvorticity takes place aloft. The converse of this is in thecase of convergence aloft.

SUMMARY

In this chapter we expanded on the subjects of convergence, divergence, and vorticity, which were firstpresented in the AG2 manual, volume 1. Our discussion

first dealt with convergence and divergence as simplemotions. The dynamics of convergent and divergentflow was covered, along with a discussion of winddirectional shear and wind speed shear. Convergenceand divergence as complex motions were thenpresented. Rules of thumb on convergence anddivergence relative to surface and u pper air featu reswere covered. The last portion of the chapter dealt withvorticity. Definitions of relative vorticity and absolutevorticity were covered. Vorticity effects on weatherprocesses was the last topic of discussion.

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CHAP TER 2

FORE CASTING UP P ER AIRSYSTEMS

To prepare surface and upper air prognostic charts,

we must first make predictions of the weather systemsfor these charts. Inasmuch as the current surface andupper air chart s reveal the curr ent stat e of the weather,

so should the prognostic charts accurately reveal thefuture stat e of the weather.

Preparing upper air and surface prognostic chartsdictates that the Aerographers Mate first begin with theupper levels and t hen t ran slate the pr ognosis downwardto the surface. The two are so interrelated that

consideration of the elements on one should not be madeindependently of the other.

Prognostic charts are constructed at the FleetNumerical Meteorology and Oceanography Center(FNMOC). The resultant products are transmitted over

their respective facsimile networks.

Overseas Meteoro logy and Oceanography(METOC) units also construct a nd t ran smit pr ognostic

charts. We are all too often inclined to take theseproducts at face value. Since these prognostic charts aregenerally for large areas, this practice could lead to an

erroneous forecast.

It is importan t th at you, the Aerograph ers Mate, not

only understand the methods by which prognostic charts

are constructed, but you should also understand theirlimitations as well. In this chapter we will discuss some

of the more common methods and rules for forecastingupper air features. In the following chapter, methodsan d techniques for pr ogging upper a ir chart s will beconsidered. These methods can be used in constructingyour own prognostic charts where data are not available

and/or to check on the prognostic charts made by othersources.

Before you read this chapter, you may find itbeneficial to review the AG2 TRAMAN, NAVEDTRA10370, volume 1,

analysis concepts.

unit 8, which discusses upper air

G E N E R A L P R O G N O S T I CC O N S I D E R AT I O N S

LEARNING OBJECTIVES: Evaluate featureson upper level char ts, and be familiar with t hevarious m eteorological pr oducts available t othe forecaster in prepar ing upper levelprognostic charts.

The forecaster must consider a l l appl icableforecasting rules, draw upon experience, and consult allavailable objective aids to pr oduce the best possibleforecast from available data.

Forecasters should examine all aspects of theweather picture from both the surface and aloft beforeissuing their forecasts. Some conditions a re deem edless important , whi le others are emphasized.Forecasters must depend heavily upon their knowledgeand experience as similar conditions yield similarconsequences. Some forecasters may decide to discarda parameter, such as surface pressure, because throughtheir experience, or the experience of others, they maydecide that it is not a decisive factor.

An objective system of forecasting certainatmospheric parameters may often exceed the skill of anexperienced forecaster. However, the objective processshould not necessarily y take precedence over a subjectivemethod, but rather the two should be used together toarrive at the most accurate forecast.

HAND D RAWN ANALYSIS

Methods and procedures u sed in th e an alysis of upper a ir charts were covered in the AG2 TRAMAN,volume 1. Accurately drawn analyses provide theforecaster with the most important tool in constructingan upper air prognostic chart. Such information aswindspeed and direction, temperature, dew pointdepression, and heights are readily available for theforecaster to integra te in to an y objective met hod forproducing a prognostic chart.

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C O M P U T E R P R O D U C TS

FNMOC provides a large number of charts fordissemination to shore and fleet units. These includea n a l y s i s a n d p r o g n o s t i c c h a r t s r a n g i n g f r o msubsu rface ocean ogra phic cha rt s t o depictions of thetroposphere, as well as a number of specializedcharts. A complete listing of the these charts iscontained in The Numerical Environmental Products

Manual , volume III ( Env i ronmen t a l P r od u c t s ) ,FLENUMMETOCCENINST 3145.2.

A P P L I C AT I O N O F S AT E L L I T E I M A G E RY

As a further aid, satellite imagery can also be usedin preparing prognostic charts. The availability of useful satellite data will vary with time and area.

O B J E C T I V E F O R E C AS T I N GT E C H N I Q U E S

LEARNING OBJECTIVES: Evaluate variousobjective forecasting techniques, includinge x t r a p o l a t i o n a n d i s o t h e r m - c o n t o u rrelationships for the movement of troughs andridges. Forecast intensity of troughs and ridges.Forecast the movement of upper level features.Forecast the intensity of upper level andassociated surface features. Lastly, forecast theformation of upper level and associated surfacefeatures.

Experience in itself is not a lways enough to forecastthe movement and/or intensity of upper air systems, but,couple the forecasters experience with basic objectivetechniques and a more accurate product will beprepared.

F O R E C A S T I N G T H E M O V E M E N T O FT R O U G H S A N D R I D G E S

Techniques covered in this section apply primarilyto long waves. Some of the techniques wil l be

applicable to short waves as well. A long wa ve is bydefinition a wave in t he m ajor belt of westerlies, whichis characterized by large length and significantamplitude. (See th e AG2 TRAMAN, volume 1, for adiscussion of long and short waves.) Therefore, the firststep in pr ogging the movement a nd int ensity of longwaves is to determine th eir limits. There ar e severa lbasic approaches to the progging of both long and

short waves. Chief ly, these are extrapolat ion,isoth erm-contour relationship, an d th e location of the jetmaximum in relation to the current in which it lies.

Ext r apo la t i on

The past history of systems affecting an area of interest is fundamental to the success of forecasting.Atmospheric systems usually change slowly, but,

continuously with time. That is, there is continuity inthe weather patterns on a sequence of weather charts.When a particular pressure system or h eight centerexhibits a tendency to continue without much change, itis said to be persistent. These concepts of persistenceand continuity ar e fundamental forecast aids.

The extrapolation procedures used in forecastingmay vary from simple extrapolation to the use of morecomplex mathematical equations and analog methodsbased on theory. The forecaster should extrapolate pastand present conditions to obtain future conditions.Extrapolation is the simplest method of forecasting bothlong an d short wave movement.

Simple extrapolation is merely the movement of thetrough or ridge to a future position based on past andcurrent movement and expected trends. It is based onthe assumption that the changes in speed of movementand intensity are slow and gradual. However, it shouldbe noted that developments frequently occur that are notrevealed from present or past indications. However, if such developments can be forecast by other techniques,allowances can be made.

Extrapolation for short periods on short waves isgenera l ly va l id . T h e m a j o r d i s a d v a n t a g e o f extrapolating the long period movement of short wavesor long waves is that past and pr esent tr ends do notcont inue indefinitely. This can be seen when weconsider a wave with a history of ret rogression. Theretr ogression will not cont inue indefinitely, and we mu stlook for indications of its reversal; that is, progressivemovement.

I so the rm-Con tou r Re l a t i onsh ips

The forecaster should always examine the longwaves for the isotherm-contour relationships, and thenapply the rules for the movement of long waves. Theserules are covered in the AG2 TRAMAN, volume 1.These rules are indicators only, but if they confirm orparallel other applied techniques, they have served theirpurpose. A number of observations and rules are statedregarding the progression, stationary characteristics, or

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retrogression of long waves. These rules are discussedin t he following t ext.

P R O G R E S S I O N O F L O N G WAVE S . Progression (eastward movement) of long waves isusually found in association with relatively short wavelengths and well defined major troughs and ridges. Atthe surface, there are usually only one or two prominentcyclones associated with each major trough aloft.Beneath the forward portion of each major ridge there

is usually a well developed surface anticyclone movingtoward the east or southeast. The 24-hour heightchanges at upper levels usually have a one-to-oneassociation with major troughs and ridges; that is,motion of maximum height fall and rise ar eas a ssociatedwith major trough and ridge motion. The tracks of theheight change centers depend on t he movement andchanges in intensity of the long waves.

STATIONARY LONG WAVE PATTERNS.Once established, stationary long wave patterns usuallypersist for a number of days. The upper airflow

associated with the long wave pat tern constitu tes asteering pattern for the smaller scale disturbances (short waves). These short waves, with their associated heightchange patterns and weak surface systems, move alongin the flow of the large scale, long wave pattern. Shortwave troughs deepen as they move through the troughsof th e long waves, and fill as t hey move through t heridges of the long waves. The same changes in intensityoccur in sea level troughs or pressure centers that areassociated with minor troughs aloft. Partly as a resultof the presence of these smaller scale systems, thetroughs a nd r idges of the sta tionary long waves ar e often

spread out and hard to locate exactly.

RETROGRESSION OF LONG WAVES. Acontinuous retrogression of long wave troughs is a rareevent. The usual type of retrogression takes place in adiscontinuous manner; a major trough weakens,accelerat es eastwar d, and becomes a minor tr ough,while a major wave trough forms to the west of theform er position of the old long wave t rough. New m ajortroughs are generally formed by the deepening of minortroughs into deep, cold troughs.

Retrogression is seldom a localized phenomenon,but appears to occur as a series of retrogressions inseveral long waves. Retrogression generally begins ina quasi-stationary long wave train when the stationarywavelength shows a significant decrease. This canhappen as a result of a decrease in zonal wind speed, orof a southward shift in the zonal westerlies. Somecharacteristics of retrogression are as follows:

l Trajectories of 24-hour height chan ge patter ns a t500-hPa deviate from th e band of maximum wind.

l New centers appear, or existing ones rapidlyincrease in intensity.

. Rapid intensification of surface cyclones occursto the west of existing major trough positions.

Loca t ion Of The Je t S t r eam

Th e AG2 TRAMAN, volume 1, discusses themigration of the jet stream both northward andsouthward. Some general considerations can be madeconcerning t his migrat ion and th e movement of wavesin the troposphere:

. In a northward migrating jet stream, a west windmaximum emerges from the tropics and graduallymoves through the lower midlatitudes. Anothermaximum, initially located in the upper midlatitudes,advan ces towar d th e Arctic Circle while weakenin g.Open progressive wave patterns with pronouncedamplitude and a decrease in the number of waves due tocutoff centers exist. The jet is well organized andtroughs extend into low latitudes.

. As the jet progresses northward, the amplitude of the long waves decrease and the cutoff lows south of thewesterlies dissipate. By the time t he jet rea ches t hemidlatitudes, a classical high zonal index (AG2TRAMAN, volume 1) situation exists. Too, we haveweak, long waves of large wavelength and smallamplitude, slowly progressive or stationary. Fewextensions of troughs into the low latitudes are present,and in this situation, the jet stream is weak anddisorganized.

. As th e jet proceeds fart her northwar d, there willoften be a sha rp br eak of high zonal index with ra pidlyincreasing wave ampli tudes a lof t . Long wavesretrograde. As the jet reaches the upper midlatitudes andinto the sub-Arctic region, it is s till the dominant featu re,while a new jet of the westerlies gradually begins toform in the subtropical regions. Long waves now begin

to increase in number, and there is a reappearance of troughs in th e tr opics. The cycle then begins a gain.

With a southward migrating jet, the processes arereversed from that of the northward moving jet. Itshould be noted that shortwaves are associated with the

jet maximum and move with about the same speed asthese jet maximums.

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F O R E C A S T I N G T H E I N T E N S I T Y O FT R O U G H S A N D R I D G E S

Forecasting the intensity of long wave troughs andridges often yields nothing more than an indication of the expected intensity; that is, greater than or less th anpresent intensity. For instance, if deepening or falling isindicated, but the extent of deepening or tilling is notdefinite, the forecaster is forced to rely on experience

and intuition in order to arrive at the amount of deepening or tilling. FNMOC upper level chartsforecast th e intensity of upper waves with a great dealof success. If available, you should check yourintensity and movement predictions against theseprognoses.

Ext rapo la t ion

Patterns on upper level charts are more persistentthan those on the surface. Therefore, extrapolationgives better results on the u pper air charts than onsurface charts. When you use height changes aloft, theprocedure is to extrapolate height change and add orsubtract th e change to the current height values.

Use of Time Differ ent ia ls

The time differential chart is discussed in t he AG2TRAMAN, volume 1.

The time differential chart constructed forth e 500-hPa level shows th e history of changes th athave taken place at the 500-hPa level at 24-hour

intervals. In considering the information on thetime differential chart, those centers with a welldefined history of movement will be of greatest value.Take into consideration not only the amount of movement, but also the changes in intensity of thecenters. Centers with no history should be treated withcaution, especially with regard to their direction of movement which is usually downstream from thecurrent position. Information derived from the timedifferential chart should be used to supplementinformation obtained from previous considerations, andwhen in agreement, used as a guide for the amount of change.

Normally, the 24-hour height rise areas can bemoved with the speed of the associated short waveridges, and the speed of the fall centers with the speedof the associated short wave troughs. It must beremembered that height change centers m ay be presentdue to convergence or divergence factors and may nothave an associated short wave trough or ridge. Be

cautious not to move a height change center with thecontour flow if it is due primarily to convergence ordivergence. However, with short wave indications, achange center will appear and move in the direction of the contour flow.

Once you have progged the movement of the heightchange centers and determined their magnitude, applythe change indicated to the height on the current 500-mbchart. You should use these points as guides inconstructing prognostic contours.

I so the rm-Con tou r Re la t ionsh ip

In long waves, deepening of troughs is associatedwith cold air advection on the west side of the troughand filling of troughs with warm air advection on thewest side of the trough. The converse is true for ridges.Warm air advection on the western side of a ridgeindicates intensification, and cold air advectionindicat es weakening. This rule is least applicable

immediate yeast of the Continental Divide in the UnitedStates, and probably east of any high mountain rangewhere westerly winds prevail aloft. In short waves,deepening of troughs is associated with cold airadvection on the west side of the trough and falling of troughs with warm air advection, particularly if a jetmaximum is in th e north erly curr ent of the t rough andtilling is indicated by warm a ir advection on th e westernside.

In reference to the above paragraph, the advectionis not the cause of the intensity changes, but rather is a

signof w h a t i s occu r r i n g. H igh l eve l

convergence/divergence is the cause.

Effec t o f Supe r Grad ien t Winds

Figure 2-1, views (A) through (D), shows the effectof the location of maximum winds on t he int ensity of troughs and ridges.

Explanation of figure 2-1 is as follows:

l When the strongest winds aloft are the westerlieson the western side of the trough, the trough deepens[fig. 2-1, view (A)].

. When the strongest winds aloft are the westerliesat the base of the trough, the trough moves rapidlyeastward and does not change in intensity [fig. 2-1, view(B)].

. When the st rongest winds ar e on t he east side of the t rough, the tr ough fills [fig. 2-1, view (C)].

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Figu re 2-1.-E ff ec t o f su p e r g r a d i e n t w i n d s o n t h e d e e p e n i n gan d f i ll ing of t rou ghs . (A) St ronges t w inds on the west s ideof t rough;(B) s t ronges t winds in south ern por t ion of t rough;(C) s t rongest winds on eas t s ide of t rough; (D) excess ivecon tou r g rad ien t s.

l Sharply curved ridges with excessive contourgradients are unstable and rotate rapidly clockwise,

causing large height rises and filling in the trough areadownstream, and large height falls in the left side of thestrong gradient ridge [fig. 2-1, view (D)].

Convergence and Dive rgence Above500 Mi ll iba r s

Study the 300-mb (or 200-mb) chart to determineareas of convergence and divergence. Note these areasof convergence and divergence.

Convergence and divergence are covered in chapter

1 of this TRAMAN, and also in the AG2 TRAMAN,volume 1. As a review of the effects of convergence anddivergence, and the changes in intensity of troughs andridges, we have the following rules:

Refer to chapter 1 for illustrations of these rules.

. Divergence and upper height falls are associatedwith high-speed winds appr oaching cyclonically curved

weak contour gradients. Divergence results in heightfalls to the left of the high-speed current.

. Convergence and upper height r ises areassociated with low-speed winds approaching straightor cyclonically curved strong contour gradients and withhigh-speed winds approaching anticyclonically curvedweak contour gradients.

F O R E C AS T I N G T H E M O VE M E N T O FU P P E R L E VE L F E ATU R E S

The movement of upper level features is discussedin t he following text .

Movemen t o f H ighs

A r e a s o f h i g h p r e s s u r e p o s s e s s c e r t a i ncharacteristics and traits. The following text discussesthese indicators for areas of high pressure.

S E M I P E R M A N E N T H I G H S . T h e

semiperma nent , subtropical highs ar e ordina rily notsubject to much day-to-day movement. When asubtropical high begins to move, it will move with thespeed and in the direction of the associated long waveridge. The movement of the long wave ridge has alreadybeen discussed. Also, seasonal movement, thoughslower and over a longer period of time, should beconsidered. These highs tend to move poleward a ndintensify in the sum mer, an d move equat orward an ddecrease in intensity in the winter.

BLOCKS. Blocks will ordinar ily persist in th e

same geographic location. Movement of blocks will bein t he direction of the st rongest winds; for example,eastward when the westerlies are strongest, andwestward when the easterlies are strongest. The speedof movement of these systems can usually bedetermined more accurately by extrapolat ion,Extrapolation should be used in moving the highs underany circumst ance, and t he r esults of this extrapolationshould be considered along with any other indications.

Some indications of intensity changes that areexhibited by lower tropospheric charts (700-500 hPa)

are as follows:l Intensification will occur with warm air

advection on the west side; weakening and decay willoccur with cold air advection on the west side; and thereis little or no change in the intensity if the isotherms aresymmetr ic with th e contour s. This low tr oposphericadvection is not the cause of the intensity change but isonly a indicator. The cause is at higher levels; for

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example, intensification is caused by high-level coldadvection and/or mass convergence.

l Under low zonal index situations, a blockinghigh will normally exist a t a northern latitude a nd willhave a pronounced effect on the systems in that area; ingeneral it will slow the movement

. Under high zonal index situations, there is astrong west to east component to the winds, and systemswill move rapidly.

M o v e m e n t o f C lo s e d L o w s

The semiperma nent Icelandic and Aleut ian lowsundergo little movement. These semipermanent lowswill decrease or increase in area of coverage;occasionally split, or elongate east-west during periodsof high zonal index. North-south displacements are dueprima rily to seasonal effects. The movement of th esesemipermanent lows is der ived pr imari ly f romextrapolation.

EXTRAPOLATION. Extra polation can be usedat times to forecast both the movement and the intensityof upper closed lows. This method should be used inconjunction with other met hods to arrive at the pr edictedposition and intensity. Figure 2-2 shows some examples

of simple extrapolation of both movement and intensity.Remember, there are ma ny variations to these patt erns,and each case must be treated individually.

Figure 2-2, view (A), illustr at es a forecast in whicha low is assumed to be moving at a constant rate andfilling. Since the low has moved 300 nautical miles inthe past 24 hours, it maybe assumed that it will move300 nautical miles in the next 24-hour period. Similarly,

since the centr al height value ha s increased by 30 metersin th e past 24 hours, you would forecast t he sam e 30meter increase for the next 24 hours. While thisprocedure is very simple, it is seldom sufficientlyaccurate. It is often refined by consulting a sequence of upper air dat a t o determine a r ate of change.

This principle is illustrated in figure 2-2, view (B).By consulting the previous charts, we find the low isfilling at a r at e of 30 meters per 24 hours ; therefore, th isconstan t r at e is predicted t o continue for t he next 24hours. However, the rate of movement is decreasing ata const ant ra te of cha nge of 100 naut ical miles in 24hours. Hence, this constant rate of change of movementis then assu med to continue for t he next 24 hours, so thelow is now predicted t o move just 200 nau tical miles inthe next 24 hours.

Figur e 2-2.-Simple ext rapola t ion of the movement and In tens i ty of a c losed low on the 700-mb char t . (A) Constantf il ling (B) constan t ra te of chan ge, (C) percenta ge ra te of chan ge.

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Frequently neither of these two situations exist, andboth the change in movement and the height centerchange occur at a proportional rate. This is illustratedin figur e 2-2, view (C). From a sequen ce of cha rt s 24hours apart, it is shown that the low is filling at adecreasing rate and also moving at a decreasing rate.The height change value is 50 percent of the value 24hours previously on t he successive charts, a nd t he ra teof movement is 75 percent. We then assume this

constan t per centa ge rat e to continue for t he next 24hours , so the low is forecast to move 225 naut ical m ilesand fill only 15 meters.

Accelerations may be handled in a similar manneras the decelerations shown in figure 2-2. Also, asequence of 12-hour charts could be used in lieu of 24-hour charts to determine past tr ends.

C R I T I C A L E C C E N T R I C I T Y. When amigratory system is unusually intense, the system mayextend vertically beyond the 300-hPa level. Advectionconsiderations, contour-isotherm relationships,convergence an d divergence considera tions, an d t helocation of the jet max will yield the movement vector.These principles are a pplied in the same ma nner a s whenthe movement of long waves are determined. Theeccentricity formula may be applied to derive amovement vector, but only when a nearly straighteastward or westward movement is apparent .Migratory lows also follow the steering principle andthe mean climatological tracks. The climatologicaltracks must be used cautiously for the obvious reasons.The rise and fall centers of the time differential chartsare of great aid in determining an extrapolatedmovement vector, and extrapolation is the primarymethod by which the movement of a closed low isdetermined.

Certain cutoff lows and migratory dynamic coldlows lend themselves to movement calculation by theeccentricity formula. The conditions under which thisformu la ma y be applied ar e:

. The low must have one or more closed contours(nearly circular in shape).

. The strongest winds must be directly north orsouth of the center. The location of the max windsdetermines the direction of movement. When thestrongest winds are t he easterlies north of the low, thelow moves westward; when the strongest winds are thewesterlies south of the low, the low will move eastward.The low will also move toward the weakest divergingcyclonic gradient and parallel to the strongest current.Systems moving eastward must have a greater speed in

order to overcome convergence upstream-there isnormally convergence east of a low system.

The eccentricity formula is written:

E c = V - V - 2 C

or

2 C = V - V - E C

where

Ec is the critical eccentricity y value.

V is the wind speed south of the closed low.

V is th e wind speed nort h of the closed low.

C is the speed of the closed low (in knots).

To obtain the value of C, it is necessary t o determinethe latitude of the center of the low and the spread (indegrees latitude) between the strongest winds in the lowand the center of the low. Apply these values to table2-1 to determine the tabular value. Apply the tabular

value to the critical eccentricity formula to obtain 2C,thus C. In determining the critical eccentricity of asystem, it is necessary to interpolate both for latitude andthe spread. A negative value for C indicates westwardmovement; a posi t ive value indicates eastwardmovement.

L O C AT I O N O F T H E J E T S T R E A M . As longas a jet maximum is situated, or moves to the westernside of a low, this low will not move. When the jet centerhas r ounded the south ern per iphery of the low, and is n otfollowed by another center upstream, the low will move

rapidly and fill.Tab le 2-1.-Cr i t i ca l Eccen t r i c i ty Va lue

La t i tude Spread (degrees latitude)(degrees)

1 3 5 10 20

80 .1 .9 2.5 -- --

70 .2 1.8 4.9 19.5 80.0

60 .3 2.6 7.1 27.0 115.0

50 .4 3.3 9.1 37.0 150.0

40 .4 4.0 10.9 43.5 175.0

30 .5 4.5 12.3 50.0 200.0

20 .5 4.9 13.3 53.0 --

10 .6 5.2 14.0 56.0 --

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ISOTHERM-CONTOUR RELATIONSHIP.Little movement will occur if the isotherms and cont oursare symmetrical (no advection). Lows will intensifyand retrogress if cold air advection occurs to the westand fill and pr ogress east ward if warm air a dvectionoccurs to the west.

F O R E C A S T IN G T H E I N T E N S I T Y O FU P P E R L E V E L AN D A S S O C IATE DS U R FA C E F E AT U R E S

Many of the same considerations used in themovement of closed centers aloft may also apply toforecasting their intensity. Extrapolation and the use of time differentials aid in forecasting the change andmagnitude of increases and decreases. Again, rise andfall indications must be used in conjunction withadvection considerations, divergence indications, andother previously discussed factors.

In t en s i ty Fo recas t ing P r inc ip l e s (H ighs )

The following text discusses how atmosphericconditions affect the forecasted intensity of highpressure systems.

l Highs undergo little or no change in intensitywhen isotherms and contours are symmetrical

. Highs intensify when warm air advection occurson the west side of the high.

. Highs weaken when cold air advection occurs on

the west side of the high.

. Blocking highs u s u a l l y i n t e n s i f y d u r i n gwestward movement and weaken during eastwardmovement.

. Convergence and height rises occur in thedownstream trough when high-speed winds with astrong gradient approach low-speed winds with ananticyclonic weak gradient. This is often the casein ridges where the west side contains the high-speedwinds; the r idge int ensifies due to th is accum ulat ionof mass. This s i tuat ion has a lso been termedovershooting. This situation can be detected at the500-hPa level, but th e 300-hP a level is better suitedbecause it is the addition or removal of mass at higherlevels that determines the height of the 500-hPacontours.

. Rise and fall centers on the time differential chartindicate the changes in intensity, both sign (increasing,decreasing) and magnitude of change, if any, in

decimeters. The magnitude of the height rises or fallscan be adjusted if other indications reveal that a slowingdown or a speeding up of the processes is occurring, andexpected to continue.

In t en s i ty Fo recas t ing P r inc ip l e s (Lows)

The following text discusses how atmosphericcondi t ions a ffec t the fo recas ted in tens i ty o f low-pressure systems.

l Lows and cutoff lows deepen when cold airadvection occurs on the west side of the trough.

. Lows fill when warm air advection occurs on thewest side of the low.

. Lows fill when a jet maximum rounds thesouthern periphery of the low.

. Lows fill when th e jet m aximum is on t he eastside of the low, if another jet max does not follow.

. Lows deepen when the jet max remains on thewest side of the low, provided the jet max to the west of the low is not preceded by another on the southernperiphery or eastern periphery of the low, for thisindicates no change in intensity.

l The 24-hour r ise and fal l centers a id inextrapolating both the change and the magnitude of fallsin moving lows. Again, these rise and fall indicationsmust be considered along with advection factors,divergence indications, and the indications of thecontour-isotherm relationships.

F O R E C AS T I N G T H E F O R M ATI O N O FU P P E R L E V E L A N D A S S O C I AT E DS U R FA C E F E AT U R E S

The following text deals with the formation of upperlevel and associated surface features , and howatmospheric features affect them.

Forma t ion Forecas t ing P r inc ip l e s (H ighs )

The following are atmospheric condition indicatorsthat are relevant to the formation of highs.

l Cold air masses of polar and Arctic origingenerally give no indication of the formation of highs atthe 500-hPa level or higher, as these airmasses normallydo not extend to this level.

. The shallow anticyclones of polar or Arcticorigin give indications of their genesis primarily on the

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surface and the 850-mb charts. The area of genesis willshow progressively colder temperatures at the surfaceand aloft; however, the drop in t he 850-mb temper atu resdoes not occur at the same rate as at the surface. This isan indication that a very strong inversion is in theprocess of form ing. The air in th e source region m ustbe relatively stagnant.

. High-level anticyclogenesis is indicated whenlow-level warm air advection is accompanied bystratospheric cold air advection. This situation hasprimary application to the formation of blocks, ashigh-level anticyclogenesis is primarily associated withth e form at ion of blocks a nd t he int ensificat ion of theridges of the subtropical highs.

. Blocks should normally be forecast to form onlyover the eastern portion of the oceans in the middle andhigh latitudes. Warm air is normally present to the northand n orthwest.

Forma t ion Forecas t ing P r inc ip l e s (Lows)

There are certain conditions required in theatm osphere, as well as certa in a tmospheric indicators,for cyclogenesis t o occur . The grea ter th e nu mber of these indicators/conditions in agreement, the greater thesuccess in forecasting cyclogenesis. Some of them arelisted below:

l An area of divergence exists aloft.

. A jet maximum on the west side of a lowindicates deepening and southward movement.

. Cold air advection in the lower troposphere andwarming in the lower stratosphere is associated with theformation of or deepening of lows.

Forma t ion Forecas t ing P r inc ip l e s(Cutoff Lows)

Another task in forecasting is that of the formationof cutoff lows. Some of the indicators are as follows:

l They genera lly form only off th e south wester ncoast of the United States and the northwestern coast of Africa.

. The upstream ridge intensifies greatly. Thisintensifying u pstream ridge conta ins a n increasing,strong, southwesterly flow.

. Strong northerlies on the west side of the trough.

. Height falls move south or southeastward.

. Strong cold air advection occurs on the west sideof the upper trough.

Cons t ruc t ing Upper Leve l P rognos t i c Cha r t s

The constant pressure prognostic chart is about totake form. The forecasted position of the long wavetroughs and ridges have been determined and depictedon the t ent at ive prognostic chart . The position of the

highs, lows, and cutoff centers were then determinedand depicted on the tentative prognostic chart. Shortwaves were treated in a similar fashion. Contours arethen depicted. The height values of the contours aredetermined by actual changes in intensity of thesystems. The pattern of the contours is largelydetermined by the position of the long waves, shortwaves, and closed pressure systems. Contours aredrawn in accordance with the following eight steps:

1. Outline the areas of warm and cold advection inthe stratum between 500 and 200 hPa, and move thethickness lines at approximately 50 percent of theindicated thickness gradient in the direction of thethermal wind.

2. Tentatively note, at several points on t he chart ,the areas of height changes on the constant pressuresurface above the existing height values.

3. Move the areas of 24-hour

us navy course navedtra 14010 - aerographer's mate 1 & c

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