introduction to descriptive physical oceanography -...

C H A P T E R

1

Introduction to Descriptive PhysicalOceanography

Oceanography is the general name given tothe scientific study of the oceans. It is histori-cally divided into physical, biological, chemical,and geological oceanography. This book is con-cerned primarily with the physics of the ocean,approached mainly, but not exclusively, fromobservations, and focusing mainly, but notexclusively, on the larger space and timescalesof the open ocean rather than on the near-coastaland shoreline regions.

Descriptive physical oceanography approachesthe ocean through both observations andcomplex numerical model output used todescribe the fluid motions as quantitatively aspossible.Dynamical physical oceanography seeksto understand the processes that govern the fluidmotions in the ocean mainly through theoreticalstudies and process-based numerical modelexperiments. This book is mainly concernedwith description based in observations (similarto previous editions of this text); however, inthis edition we include some of the concepts ofdynamical physical oceanography as an impor-tant context for the description. A full treatmentof dynamical oceanography is contained in othertexts. Thermodynamics also clearly enters into ourdiscussion of the ocean through the processesthat govern its heat and salt content, and there-fore its density distribution.

Chapter 2 describes the ocean basins and theirtopography.Thenext three chapters introduce thephysical (and some chemical) properties of fresh-water and seawater (Chapter 3), an overview ofthe distribution of water characteristics (Chapter 4),and the sources and sinks of heat and freshwater(Chapter 5). The next three chapters cover datacollection and analysis techniques (Chapter 6and supplemental material listed as Chapter S6on the textbook Web site http://booksite.academicpress.com/DPO/; “S” denotes supple-mentalmaterial.), an introduction to geophysicalfluid dynamics for graduate students who havevarying mathematics backgrounds (Chapter 7),and then basicwaves and tideswith an introduc-tion to coastal oceanography (Chapter 8). Thelast six chapters of the book introduce the circu-lation and water properties of each of the indi-vidual oceans (Chapters 9 through 13) endingwith a summary of the global ocean in Chapter 14.

Accompanying the text is the Web sitementioned in the previous paragraph. It hasfour aspects:

1. Textbook chapters on climate variability andoceanographic instrumentation that do notappear in the print version

2. Expanded material and additional figures formany other chapters

1Descriptive Physical Oceanography

� 2011. Lynne Talley, George Pickard, William Emery and James Swift.

Published by Elsevier Ltd. All rights reserved.

3. A full set of tutorials for descriptiveoceanography with data and sample scriptsprovided based on the Java Ocean Atlas(Osborne & Swift, 2009)

4. All figures from the text, with many more incolor than in the text, for lectures andpresentations.

1.1. OVERVIEW

There are many reasons for developing ourknowledge of the ocean. Near-shore currentsand waves affect navigation and construction ofpiers, breakwaters, and other coastal structures.The large heat capacity of the oceans exertsa significant and in some cases a controlling effecton the earth’s climate. The ocean and atmosphereinteract on short to long timescales; for example,the El NinoeSouthern Oscillation (ENSO)phenomenon that, although driven locally in thetropical Pacific, affects climate on timescales ofseveral years over much of the world. To under-stand these interactions it is necessary to under-stand the coupled ocean-atmosphere system. Tounderstand the coupled system, it is first neces-sary to have a solid base of knowledge aboutboth the ocean and atmosphere separately.

In these and many other applications, knowl-edge of the ocean’s motion and water propertiesis essential. This includes the major oceancurrents that circulate continuously but withfluctuating velocity and position, the variablecoastal currents, the rise and fall of the tides,and the waves generated by winds or earth-quakes. Temperature and salt content determinedensity and hence vertical movement. They alsoaffect horizontal movement as the density affectsthe horizontal pressure distribution. Sea ice hasits own full set of processes and is important fornavigation, ocean circulation, and climate. Otherdissolved substances such as oxygen, nutrients,and other chemical species, and even some ofthe biological aspects such as chlorophyllcontent, are used in the study of ocean physics.

Our present knowledge in physical oceanog-raphy rests on an accumulation of data, most ofwhich were gathered during the past 150 years,with a large increase of in situ data collection(within the actual water) starting in the 1950sand an order of magnitude growth in availabledata as satellites began making ocean measure-ments (starting in the 1970s).

A brief history of physical oceanographywithillustrations is provided as supplemental mate-rial on the textbook Web site (Chapter 1 supple-ment is listed as Chapter S1 on the Web site).Historically, sailors have always been concernedwith how ocean currents affect their ships’courses as well as changes in ocean temperatureand surface condition. Many of the earlier navi-gators, such as Cook and Vancouver, made valu-able scientific observations during their voyagesin the late 1700s, but it is generally consideredthat Mathew Fontaine Maury (1855) started thesystematic large-scale collection of ocean currentdata using ship’s navigation logs as his source ofinformation. The firstmajor expedition designedexpressly to study all scientific aspects of theoceans was carried out on the British H.M.S.Challenger that circumnavigated the globe from1872 to 1876. The first large-scale expeditionorganized primarily to gather physical oceano-graphic data was carried out on the German FSMeteor, which studied the Atlantic Ocean from1925 to 1927 (Spiess, 1928). A number of photosfrom that expedition are reproduced on theaccompanying Web site. Some of the earliesttheoretical studies of the sea included the surfacetides by Newton, Laplace, and Legendre (e.g.,Wilson, 2002) and waves by Gerstner and Stokes(e.g., Craik, 2005). Around 1896, some of theScandinavian meteorologists started to turntheir attention to the ocean, because dynamicalmeteorology and dynamical oceanographyhave many common characteristics. Currentknowledge of dynamical oceanography owesits progress to the work of Bjerknes, Bjerknes,Solberg, and Bergeron (1933), Ekman (1905,1923), Helland-Hansen (1934), and others.

1. INTRODUCTION TO DESCRIPTIVE PHYSICAL OCEANOGRAPHY2

The post-war 1940s through 1960s began toproduce much of the data and especially theoret-ical understanding for large-scale ocean circula-tion. With the advent of moored and satelliteinstrumentation in the 1960s and 1970s, thesmaller scale, energetically varying part of theocean circulation d the mesoscale d began tobe studied in earnest. Platforms expanded fromresearch and merchant ships to global satelliteand autonomous instrument sampling. Futuredecades should take global description andmodeling to even smaller scales (submesoscale)as satellite observations and numerical modelingresolution continue to evolve, and different typesof autonomous sampling within the watercolumn become routine. Physical oceanographyhas retained an aspect of individual explorationbut large, multi-investigator, multinationalprograms have increasingly provided many ofthe new data sets and understanding. Currentresearch efforts in physical oceanography arefocused on developing an understanding of thevariability of the ocean and its relation to theatmosphere and climate as well as continuing todescribe its steady-state conditions.

1.2. SPACE AND TIMESCALES OFPHYSICAL OCEANOGRAPHIC

PHENOMENA

The ocean is a fluid in constant motion witha very large range of spatial and temporalscales. The complexity of this fluid is nicely rep-resented in the sea surface temperature image ofthe Gulf Stream captured from a satellite shownin Figure 1.1a. The Gulf Stream is the westernboundary current of the permanent, large-scaleclockwise gyre circulation of the subtropicalNorth Atlantic. The Gulf Stream has a width of100 km, and its gyre has a spatial scale of thou-sands of kilometers. The narrow, warm core ofthe Gulf Stream (red in Figure 1.1a) carrieswarm subtropical water northward from theCaribbean, loops through the Gulf of Mexico

around Florida and northward along the eastcoast of North America, leaves the coast atCape Hatteras, and moves out to sea. Itsstrength and temperature contrast decay east-ward. Its large meanders and rings, with spatialscales of approximately 100 km, are consideredmesoscale (eddy) variability evolving on time-scales of weeks. The satellite image also showsthe general decrease in surface temperaturetoward the north and a large amount of small-scale eddy variability. The permanence of theGulf Stream is apparent when currents andtemperatures are averaged in time. Averagingmakes the Gulf Stream appear wider, especiallyafter the separation at Cape Hatteras where thewide envelope of meanders creates a wide,weak average eastward flow.

The Gulf Stream has been known and chartedfor centuries, beginning with the Spanish expe-ditions in the sixteenth century (e.g., Peterson,Stramma, & Kortum 1996). It was first mappedaccurately in 1769 by Benjamin Franklin

FIGURE 1.1 (a) Sea surface temperature from a satelliteadvanced very high resolution radiometer (AVHRR)instrument (Otis Brown, personal communication, 2009).This figure can also be found in the color insert.

SPACE AND TIMESCALES OF PHYSICAL OCEANOGRAPHIC PHENOMENA 3

working together with whaling captain TimothyFolger (Figure 1.1b; from Richardson, 1980a).1

The narrow current along the coast of theUnited States is remarkably accurate. Thewidening envelope of the Franklin/Folgercurrent after separation from Cape Hatteras isan accurate depiction of the envelope ofmeandering apparent in the satellite image.

When time-mean averages of the Gulf Streambased on modern measurements are con-structed, they look remarkably similar to thisFranklin/Folger map.

The space and timescales of many of theimportant physical oceanography processes areshown schematically in Figure 1.2. At the small-est scale is molecular mixing. At small,

FIGURE 1.1 (b) Franklin-Folger map of the Gulf Stream. Source: From Richardson (1980a).

1 Franklin noted on his frequent trips between the United States and Europe that some trips were considerably quicker than

others. He decided that this was due to a strong ocean current flowing from the west to the east. He observed marked

changes in surface conditions and reasoned that this ocean current might be marked by a change in sea surface

temperature. Franklin began making measurements of the ocean surface temperature during his travels. Using a simple

mercury-in-glass thermometer, he was able to determine the position of the current.


macroscopic scales of centimeters, microstruc-ture (vertical layering at the centimeter level)and capillary waves occur. At the slightly largerscale of meters surface waves are found, whichhave rapid timescales and somewhat long-livedvertical layers (fine structure). At scales of tensof meters are the internal waves with timescalesof up to a day. Tides have the same timescalesas internal waves, but much larger spatial scalesof hundreds to thousands of kilometers. Surfacewaves, internal waves, and tides are describedin Chapter 8.

Mesoscale eddies and strong ocean currentssuch as the Gulf Stream are found at spatialscales of tens of kilometers to several hundredkilometers and timescales of weeks to years(Figure 1.1a,b). The large-scale ocean circulationhas a spatial scale the size of ocean basins up tothe global ocean and a timescale ranging fromseasonal to permanent, which is the timescaleof plate tectonics that rearranges the oceanboundaries (Chapter 2). The timescales forwind-driven and thermohaline circulation inFigure 1.2 are actually the same for circulationof the flow through those systems (ten yearsaround the gyre, hundreds of years through thefull ocean); these are time-mean features of the

ocean and havemuch longer timescales. Climatevariability affects the ocean, represented inFigure 1.2 by the El Nino, which has an interan-nual timescale (several years; Chapter 10);decadal and longer timescales of variability ofthe ocean circulation and properties are alsoimportant and described in each of the oceanbasin and global circulation chapters.

We see in Figure 1.2 that short spatial scalesgenerally have short timescales, and long spatialscales generally have long timescales. There aresome exceptions to this, most notably in thetides and tsunamis as well as in some fine-struc-ture phenomena with longer timescales thanmight be expected from their short spatialscales.

In Chapter 7, where ocean dynamics are dis-cussed, some formal non-dimensional parametersincorporating the approximate space and time-scales for these different types of phenomenaare introduced (see also Pedlosky, 1987). Anon-dimensional parameter is the ratio ofdimensional parameters with identical dimen-sional scales, such as time, length, mass, etc.,which are intrinsic properties of the flowphenomenon being described or modeled. Ofspecial importance is whether the timescale of

Bubbles

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FIGURE 1.2 Time and spacescales of physical oceanographicphenomena from bubbles andcapillary waves to changes inocean circulation associated withEarth’s orbit variations.

SPACE AND TIMESCALES OF PHYSICAL OCEANOGRAPHIC PHENOMENA 5

an oceanmotion is greater than or less than abouta day, which is the timescale for the earth’s rota-tion. Earth’s rotation has an enormous effect onhow the ocean water moves in response toa force; if the force and motion are sustainedfor days or longer, then the motion is stronglyinfluenced by the rotation. Therefore, an espe-cially useful parameter is the ratio of the time-scale of Earth’s rotation to the timescale of themotion. This ratio is called the Rossby number.For the very small, fast motions in Figure 1.2,this ratio is large and rotation is not important.For the slow, large-scale part of the spectrum,the Rossby number is small and Earth’s rotationis fundamental. A second very important non-dimensional parameter is the ratio of the verticallength scale (height) to the horizontal length

scale; this is called the aspect ratio. For large-scaleflows, this ratio is very small since the verticalscale can be no larger than the ocean depth. Forsurface and internal gravity waves, the aspectratio is order 1. We will also see that dissipationis very weak in the sense that the timescale fordissipation to act is long compared with boththe timescale of Earth’s rotation and the time-scale for the circulation to move water fromone place to another; the relevant non-dimen-sional parameters are the Ekman number andReynolds number, respectively. Understandinghow the small Rossby number, small aspectratio, and nearly frictionless fluid ocean behaveshas depended on observations of the circulationandwater propertiesmade over the past century.These are the principal subjects of this text.


C H A P T E R

2

Ocean Dimensions, Shapes,and Bottom Materials

2.1. DIMENSIONS

The oceans are basins in the surface of thesolid earth containing salt water. This chapterintroduces some nomenclature and directsattention to features of the basins that havea close connection with the ocean’s circulationand dynamical processes that are of importanceto the physical oceanographer. More detaileddescriptions of the geology and geophysics ofthe ocean basins are given in Seibold and Berger(1982), Kennett (1982), Garrison (2001), andThurman and Trujillo (2002), among others.Updated data sets, maps, and information areavailable from Web sites of the NationalGeophysical Data Center (NGDC) of theNational Oceanic and Atmospheric Administra-tion (NOAA) and from the U.S. GeologicalSurvey (USGS).

The major ocean areas are the Atlantic Ocean,the Pacific Ocean, the Indian Ocean, the ArcticOcean, and the Southern Ocean (Figure 2.1).The first four are clearly divided from eachother by land masses, but the divisions betweenthe Southern Ocean and oceans to its north aredetermined only by the characteristics of theocean waters and their circulations. Thegeographical peculiarities of each ocean aredescribed in Section 2.11.

The shape, depth, and geographic location ofan ocean affect the general characteristics of itscirculation. Smaller scale features, such as loca-tions of deep sills and fracture zones, seamounts,and bottom roughness, affect often importantdetails of the circulation and ofmixing processesthat are essential to forcing andwater properties.The Atlantic has a very marked “S” shape whilethe Pacific has a more oval shape. The Atlanticand Indian Oceans are roughly half the east-west width of the Pacific Ocean, which impactsthe way that each ocean’s circulation adjusts tochanges in forcing. The Indian Ocean has nohigh northern latitudes, and therefore no possi-bility of cold, dense water formation. The edgesof the Pacific are ringed with trenches, volca-noes, and earthquakes that signal the gradualdescent of the ocean bottom crustal “plates”under the surrounding continental plates. Incontrast, the Atlantic is the site of dynamicseafloor spreading as material added in thecenter of the Mid-Atlantic Ridge (MAR) pushesthe plates apart, enlarging the Atlantic Oceanby a few centimeters each year.

Marginal seas are fairly large basins of saltwater that are connected to the open ocean byone or more fairly narrow channels. Those thatare connected by very few channels are some-times called mediterranean seas after the

7Descriptive Physical Oceanography

� 2011. Lynne Talley, George Pickard, William Emery and James Swift.

Published by Elsevier Ltd. All rights reserved.

prototype, the (European) Mediterranean Sea.The Mediterranean provides an example ofa negative water balance in a sea with lessinflow (river runoff and precipitation) thanevaporation. An excellent example of a positivewater balance marginal sea (with net precipita-tion) is found in the Black Sea, which connectswith the Mediterranean Sea. Both of these seasare discussed further in Chapters 5 and 9. Otherexamples of marginal seas that are separatedfrom the open ocean bymultiple straits or islandchains are the Caribbean Sea, the Sea of Japan,the Bering Sea, the North Sea, the Baltic Sea,and so forth.

The term sea is also used for a portion of anocean that is not divided off by land but haslocal distinguishing oceanographic characteris-tics; for example the Norwegian Sea, the Labra-dor Sea, the Sargasso Sea, and the Tasman Sea.

More of the earth’s surface is covered by seathan by land, about 71% sea to 29% land. (Themost recent elevation data for the earth’s surface,used to construct Figure 2.2, show that 70.96% ofthe earth is ocean; see Becker et al., 2009.)Furthermore, the proportion of water to land inthe Southern Hemisphere is much greater (4:1)than in the Northern Hemisphere (1.5:1). Inarea, the Pacific Ocean is about as large as theAtlantic and Indian Oceans combined. If theneighboring sectors of the Southern Ocean areincluded with the three main oceans north of it,the Pacific Ocean occupies about 46% of the totalworld ocean area, the Atlantic Ocean about 23%,the Indian Ocean about 20%, and the rest,combined, about 11%.

The average depth of the oceans is close to4000 m while the marginal seas are generallyabout 1200 m deep or less. Relative to sea level,

180˚ 120˚W 60˚W 0˚ 60˚E 120˚E 180˚

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FIGURE 2.1 Map of the world based on ship soundings and satellite altimeter derived gravity at 30 arc-second reso-lution. Data from Smith & Sandwell (1997); Becker et al. (2009); and SIO (2008).

2. OCEAN DIMENSIONS, SHAPES, AND BOTTOM MATERIALS8

the oceans aremuchdeeper than the land is high.While only 11% of the land surface of the earth ismore than 2000 m above sea level, 84% of the seabottom is more than 2000 m deep. However, themaxima are similar: the height of Mt. Everest isabout 8848 m, while the maximum depth in theoceans is 11,034 m in the Mariana Trench in thewestern North Pacific. Figure 2.2 shows thedistributions of land elevations and of sea depthsrelative to sea level in 100 m intervals as thepercentage of the total area of the earth’s surface.This figure is based on the most recent globalelevation and ocean bathymetry data fromD. Sandwell (Becker et al., 2009). (It is similar toFigure 2.2 using 1000 m bins that appeared inprevious editions of this text, based on datafrom Kossina, 1921 and Menard & Smith, 1966,

but the 100 m bins allow much more differentia-tion of topographic forms.)

Although the average depth of the oceans,4 km, is a considerable distance, it is smallcompared with the horizontal dimensions ofthe oceans, which are 5000 to 15,000 km. Rela-tive to the major dimensions of the earth, theoceans are a thin skin, but between the seasurface and the bottom of the ocean there isa great deal of detail and structure.

2.2. PLATE TECTONICS ANDDEEP-SEA TOPOGRAPHY

The most important geophysical processaffecting the shape and topography of the ocean

Maximum (Mariana Trench)

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FIGURE 2.2 Areas of Earth’s surface above and below sea level as a percentage of the total area of Earth (in 100 mintervals). Data from Becker et al. (2009).

PLATE TECTONICS AND DEEP-SEA TOPOGRAPHY 9

basins is the movement of the earth’s tectonicplates, described thoroughly in texts such asThurman and Trujillo (2002). The plate bound-aries are shown in Figure 2.3. Seafloor spreadingcreates new seafloor as the earth’s plates spreadapart. This creates the mid-ocean ridge system;the mid-ocean ridges of Figure 2.1 correspondto plate boundaries. The ocean plates spreadapart at rates of about 2 cm/year (Atlantic) to16 cm/year (Pacific), causing extrusion ofmagma into the surface at the centers of theridges. Over geologic time the orientation ofthe earth’s magnetic field has reversed, causingthe ferromagnetic components in the moltennew surface material to reverse. Spreading atthe mid-ocean ridge was proven by observingthe reversals in the magnetic orientations inthe surface material. These reversals permitdating of the seafloor (Figure 2.3). The recur-rence interval for magnetic reversals is approxi-mately 500,000 to 1,000,000 years.

The 14,000 km long MAR is a tectonicspreading center. It is connected to the globalmid-ocean ridge, which at more than 40,000 kmlong, is the most extensive feature of the earth’stopography. Starting in the Arctic Ocean, themid-ocean ridge extends through Icelanddown the middle of the Atlantic, wraps aroundthe tip of Africa, and then winds through theIndian and Pacific Oceans, ending in the Gulfof California. In all oceans, the mid-ocean ridgeand other deep ridges separate the bottomwaters, as can be seen from different waterproperties east and west of the ridge.

Deep and bottom waters can leak across theridges through narrow gaps, called fracturezones, which are lateral jogs in the spreadingcenter. The fracture zones are roughly verticalplanes, perpendicular to the ridge, on eitherside of which the crust has moved in oppositedirections perpendicular to the ridge. Thereare many fracture zones in the mid-ocean

FIGURE 2.3 Sea floor age (millions of years). Black lines indicate tectonic plate boundaries. Source: From Muller, Sdrolias,Gaina, and Roest (2008).


ridges. One example that is important asa pathway for abyssal circulation in the Atlanticis the Romanche Fracture Zone through theMAR close to the equator. Another example isthe pair of fracture zones in the South Pacific(Eltanin and Udintsev Fracture Zones,Figure 2.12) that steer the Antarctic CircumpolarCurrent (ACC).

At some edges of the tectonic plates, one platesubducts (moves under) another. Subduction isaccompanied on its landward side by volcanoesand earthquakes. Subduction creates deeptrenches that are narrow relative to their lengthand have depths to 11,000 m. The deepest partsof the oceans are in these trenches. The majorityof the deep trenches are in the Pacific: the Aleu-tian, Kurile, Tonga, Philippine, and Mariana.There are a few in other oceans such as thePuerto Rico and the South Sandwich Trenchesin the Atlantic and the Sunda Trench in theIndian Ocean. Trenches are often shaped likean arc of a circle with an island arc on one side.Examples of island arcs are the Aleutian Islands(Pacific), the Lesser Antilles (Atlantic), and theSunda Arc (Indian). The landward side ofa trench extends as much as 10,000 m from thetrench bottom to the sea surface, while the otherside is only half as high, terminating at theocean depth of about 5000 m.

Trenches can steer or impact boundarycurrents that are in deep water (Deep WesternBoundary Currents) or upper ocean boundarycurrents that are energetic enough to extend tothe ocean bottom, such as western boundarycurrents of the wind-driven circulation. Exam-ples of trenches that impact ocean circulationare the deep trench system along the westernand northern boundary of the Pacific and thedeep trench east of the Caribbean Sea in theAtlantic.

Younger parts of the ocean bottom are shal-lower than older parts. As the new seafloorcreated at seafloor spreading centers ages, itcools by losing heat into the seawater aboveand becomes denser and contracts, which

causes it to be deeper (Sclater, Parsons, & Jau-part, 1981). Ocean bottom depths range from 2to 3 km for the newest parts of the mid-oceanridges to greater than 5 km for the oldest, ascan be seen by comparing the maps of seafloorage and bathymetry (Figures 2.1 and 2.3).

The rate of seafloor spreading is so slow that ithas no impact on the climate variability that weexperience over decades to millennia, nor does itaffect anthropogenic climate change. However,over many millions of years, the geographiclayout of Earth has changed. The paleocircula-tion patterns of “deep time,” when the conti-nents were at different locations, differed fromthe present patterns; reconstruction of thesepatterns is an aspect of paleoclimate modeling.By studying and understanding present-daycirculation, we can begin to credibly model thepaleocirculation, which had the same physicalprocesses (such as those associated with theearth’s rotation, wind and thermohalineforcing, boundaries, open east-west channels,equatorial regions, etc.), but with differentocean basin shapes and bottom topography.

Ocean bottom roughness affects ocean mix-ing rates (Sections 7.2.4 and 7.3.2). The overallroughness varies by a factor of 10. Roughnessis a function of spreading rates and sedimenta-tion rates. New seafloor is rougher than oldseafloor. Slow-spreading centers producerougher topography than fast spreadingcenters. Thus the slow-spreading MAR isrougher than the faster spreading East PacificRise (EPR; Figure 2.4). Slow-spreading ridgesalso have rift valleys at the spreading center,whereas fast-spreading ridges have an elevatedridge at the spreading center. Much of theroughness on the ridges can be categorized asabyssal hills, which are the most common land-form on Earth. Abyssal hills are evident inFigures 2.4 and 2.5b, all along the wide flanksof the mid-ocean ridge.

Individual mountains (seamounts) are widelydistributed in the oceans. Seamounts stand outclearly above the background bathymetry. In


the maps in Figure 2.4b there are someseamounts on the upper right side of the panel.In the vertical cross-section in Figure 2.5b,seamounts are distinguished by their greaterheight compared with the abyssal hills. Theaverage height of seamounts is 2 km.Seamounts that reach the sea surface formislands. A guyot is a seamount that reachedthe surface, was worn flat, and then sank againbelow the surface. Many seamounts andislands were created by volcanic hotspotsbeneath the tectonic plates. The hotspots arerelatively stationary in contrast to the platesand as the plates move across the hotspots,chains of seamounts are formed. Examples

include the Hawaiian Islands/Emperor Sea-mounts chain, Polynesian island chains, theWalvis Ridge, and the Ninetyeast Ridge in theIndian Ocean.

Seamounts affect the circulation, especiallywhen they appear in groups as they do inmany regions; for instance, the Gulf Streampasses through the New England Seamounts,which affect the Gulf Stream’s position andvariability (Section 9.3). Seamount chains alsorefract tsunamis, which are ocean wavesgenerated by submarine earthquakes that reactto the ocean bottom as they propagate longdistances from the earthquake source (Section8.3.5).

FIGURE 2.4 Seafloor topography for a portion of (a) the fast-spreading EPR and (b) the slow-spreading MAR. Note theridge at the EPR spreading center and rift valley at the MAR spreading center. This figure can also be found in the colorinsert. (Sandwell, personal communication, 2009.)


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FIGURE 2.5 (a) Schematic section through ocean floor to show principal features. (b) Sample of bathymetry, measuredalong the South Pacific ship track shown in (c).


2.3. SEAFLOOR FEATURES

The continents form the major lateral bound-aries to the oceans. The detailed features of theshoreline and of the sea bottom are importantbecause of the way they affect circulation. Start-ing from the land, themain divisions recognizedare the shore, the continental shelf, the conti-nental slope and rise, and the deep-seabottom, part of which is the abyssal plain(Figure 2.5a, b). Some of the major features ofthe seafloor, including mid-ocean ridges, tren-ches, island arcs, and seamounts, are the resultof plate tectonics and undersea volcanism(Section 2.2 and Figure 2.3).

In some of the large basins the seafloor is verysmooth, possibly more so than the plains areason land. Sedimentation, which is mostly dueto the incessant rain of organic matter from theupper ocean, covers the rough bottom andproduces large regions of very smooth topog-raphy. Stretches of the abyssal plain in thewestern North Atlantic have been measured tobe smooth within 2 m over distances of 100 km.The ocean bottom in the northeast IndianOcean/Bay of Bengal slopes very smoothlyfrom 2000 m down to more than 5000 m over3000 km. This smoothness is due to sedimenta-tion from the Ganges and Brahmaputra Riversthat drain the Himalayas. Bottom sedimentscan be moved around by deep currents; forma-tion of undersea dunes and canyons is common.Erosional features in deep sediments havesometimes alerted scientists to the presence ofdeep currents.

Bottom topography often plays an importantrole in the distribution of water masses and thelocation of currents. For instance, bottom watercoming from the Weddell Sea (Antarctica) isunable to fill the eastern part of the Atlanticbasin directly due to the height of the WalvisRidge (South Atlantic Ocean). Instead, thebottom water travels to the north along thewestern boundary of the South Atlantic, findsa deep passage in the MAR, and then flows

south to fill the basin east of the ridge. At shal-lower depths the sills (shallowest part ofa channel) defining the marginal seas stronglyinfluences both the mid-level currents and thedistribution of water masses associated withthe sea. Coastal upwelling is a direct conse-quence of the shape of the coast and its relatedbottom topography. Alongshore currents areoften determined by the coastal bottom topog-raphy and the instabilities in this system candepend on the horizontal scales of the bottomtopography. Near the shore bottom topographydictates the breaking of surface gravity wavesand also directly influences the local tidalexpressions.

Much of the mixing in the ocean occurs nearthe boundaries (including the bottom). Micro-structure observations in numerous regions,and intensive experiments focused on detectionof mixing and its genesis, suggest that flow ofinternal tides over steep bottom slopes in thedeep ocean is amajor mechanism for dissipatingthe ocean’s energy. Ocean bottom slopescomputed from bathymetry collected alongship tracks show that the largest slopes tend tooccur on the flanks of the fastest spreadingmid-ocean ridges. With bathymetric slopescomputed from the most recent bathymetricdata (Figure 2.6) and information about theocean’s deep stratification, it appears the flanksof the mid-ocean ridges of the Atlantic,Southern Ocean, and Indian Ocean could bethe most vigorous dissipation sites of oceanenergy (Becker & Sandwell, 2008).

2.4. SPATIAL SCALES

Very often some of the characteristics of theocean are presented by drawing a verticalcross-section of a part of the oceans, such asthe schematic depiction of ocean floor featuresin Figure 2.5a. An illustration to true scalewould have the relative dimensions of theedge of a sheet of paper and would be either


too thin to show details or too long to be conve-nient. Therefore, we usually distort our cross-section by making the vertical scale much largerthan the horizontal one. For instance, we mightuse a scale of 1 cm to represent 100 km horizon-tally while depths might be on a scale of 1 cmto represent 100 m (i.e., 0.1 km). In this casethe vertical dimensions on our drawing wouldbe magnified 1000 times compared with thehorizontal ones (a vertical exaggeration of1000:1). This gives us room to show the detail,but it also exaggerates the slope of the seabottom or of contours of constant water pro-perties (isopleths) drawn on the cross-section(Figure 2.5b). In reality, such slopes are far less

than they appear on the cross-section drawings;for instance, a line of constant temperature(isotherm) with a real slope of 1 in 10,000 wouldappear as a slope of 1 in 10 on the plot.

2.5. SHORE, COAST, AND BEACH

The shore is defined as that part of the land-mass, close to the sea, that has been modifiedby the action of the sea. The words shore andcoast have the same meaning. Shorelines (coasts)shift over time because of motion of the landover geologic time, changes in sea level, anderosion and deposition. The sedimentary record

FIGURE 2.6 Mean slope of the ocean bottom, calculated from shipboard bathymetry and interpolated to a 0.5 degreegrid. Source: From Becker and Sandwell (2008).

SHORE, COAST, AND BEACH 15

shows a series of marine intrusions and retreatscorresponding to layers that reflect periodswhen the surface was above and below sealevel. Variations in sea level between glacialand interglacial periods have been as much as120 m. The ability of the coast to resist theerosional forces of the ocean depends directlyon the type of material that makes up the coast.Sands are easily redistributed by the oceancurrents whereas granitic coasts are slow toerode. Often sea level changes are combinedwith the hydrologic forces of an estuary, whichdramatically change the dynamical relationshipbetween the ocean and the solid surface.

The beach is a zone of unconsolidated particlesat the seaward limit of the shore and extendsroughly from the highest to the lowest tidelevels. The landward limit of the beach mightbe vegetation, permanent sand dunes, or humanconstruction. The seaward limit of a beach,where the sediment movement on- and offshoreceases, is about 10 m deep at low tide.

Coasts can be classified in many differentways. In terms of long timescales (such as thoseof plate tectonics; Section 2.2), coasts and conti-nental margins can be classified as active orpassive. Active margins, with active volcanism,faulting, and folding, are like those in much ofthe Pacific and are rising. Passive margins, likethose of the Atlantic, are being pushed in frontof spreading seafloor, are accumulating thickwedges of sediment, and are generally falling.Coasts can be referred to as erosional or deposi-tional depending on whether materials areremoved or added. At shorter timescales, wavesand tides cause erosion and deposition. Atmillennial timescales, changes in mean sea levelcause materials to be removed or added.Erosional coasts are attacked by waves andcurrents, both of which carry fine material thatabrades the coast. The waves create alongshoreand rip currents (Section 8.3) that carry theabraded material from the coastline along andout to sea. This eroded material can be joinedby sediments discharged from rivers and form

deltas. This type of erosion is fastest on high-energy coasts with large waves, and slowest onlow-energy coasts with generally weak wavefields. Erosion occurs more rapidly in weakermaterials than in harder components. Thesevariations in materials allow erosive forces tocarve characteristic features on coastlines suchas sea cliffs and sea caves, and to create an alter-nation between bays and headlands.

Beaches result when sediment, usually sand,is transported to places suitable for continueddeposition. Again these are often the quietbays between headlands and other areas oflow surf activity. Often a beach is in equilibrium;new sand is deposited to replace sand that isscoured away. Evidence for this process can beseen by how sand accumulates against newstructures built on the shore, or by how it isremoved from a beach when a breakwater isbuilt that cuts off the supply of sand beyond it.On some beaches, the sand may be removedby currents associated with high waves at oneseason of the year and replaced by differentcurrents associated with lower waves at anotherseason. These currents are influenced byseasonal and interannual wind variations.

Sea level, which strongly affects coasts, isaffected by the total amount of water in theocean, changes in the containment volume ofthe world’s ocean, and changes in the tempera-ture/salinity characteristics of the ocean thatalter its density and hence cause the water toexpand or contract. Changes in the total amountof water are due primarily to changes in thevolume of landfast ice, which is contained inice sheets and glaciers. (Because sea ice floatsin water, changes in sea ice volume, such asthat in the Arctic or Antarctic, do not affect sealevel.) Changes in containment volume aredue to tectonics, the slow rebound of continents(continuing into the present) after the melt oflandfast ice after the last deglaciation, andrebound due to the continuing melt of glacialice. Changes in heat content cause seawater toexpand (heating) or contract (cooling).


Sea level rose 20 cm from 1870 to 2003,including 3 cm in just the last 10 years(1993e2003). Because good global observationsare available for that last 10 years, it is possibleto ascribe 1.6 cm to thermal expansion, 0.4 cmto Greenland and Antarctic ice sheet melt, and0.8 cm to other glacial melt with a residual of0.3 cm. Sea level is projected to rise 30 � 10 cmin the next 100 years mainly due to warmingof the oceans, which absorb most of the anthro-pogenic heat increase in the earth’s climatesystem. (See Bindoff et al., 2007 in the 4th assess-ment report of the Intergovernmental Panel onClimate Change.)

2.6. CONTINENTAL SHELF, SLOPE,AND RISE

The continental shelf extends seaward fromthe shore with an average gradient of 1 in 500.Its outer limit (the shelf break) is set where thegradient increases to about 1 in 20 (on average)to form the continental slope down to the deepsea bottom. The shelf has an average width of65 km. In places it is much narrower than this,while in others, as in the northeastern BeringSea or the Arctic shelf off Siberia, it is as muchas ten times this width. The bottom material isdominantly sand with less common rock ormud. The shelf break is usually clearly evidentin a vertical cross-section of the sea bottomfrom the shore outward. The average depth atthe shelf break is about 130 m. Most of theworld’s fisheries are located on the continentalshelves for a multitude of reasons includingproximity of estuaries, depth of penetration ofsunlight compared with bottom depth, andupwelling of nutrient-rich waters onto someshelves, particularly those off western coasts.

The continental slope averages about 4000 mvertically from the shelf to the deep-sea bottom,but in places extends as much as 9000 m verti-cally in a relatively short horizontal distance.In general, the continental slope is considerably

steeper than the slopes from lowland to high-land on land. The material of the slope ispredominantly mud with some rock outcrops.The shelf and slope typically include submarinecanyons, which are of worldwide occurrence.These are valleys in the slope, either V-shapedor with vertical sides, and are usually foundoff coasts with rivers. Some, usually in hardgranitic rock, were originally carved as riversand then submerged, such as around the Medi-terranean and southern Baja, California. Others,commonly in softer sedimentary rock, areformed by turbidity currents described in thenext paragraph. The lower part of the slope,where it grades into the deep-sea bottom, isreferred to as the continental rise.

Turbidity currents (Figure 2.7) are common oncontinental slopes. These episodic events carrya mixture of water and sediment and are drivenby the unstable sediments rather than by forceswithin the water. In these events, material buildsup on the slope until it is no longer stable andthe force of gravity wins out. Large amounts ofsediment and bottom material crash down theslope at speeds up to 100 km/h. These eventscan snap underwater cables. The precise condi-tions that dictate when a turbidity currentoccurs vary with the slope of the valley andthe nature of the material in the valley. Turbiditycurrents carve many of the submarine canyonsfound on the slopes. Some giant rivers, such asthe Congo, carry such a dense load of sus-pended material that they form continuousdensity flows of turbid water down theircanyons.

2.7. DEEP OCEAN

From the bottom of the continental slope, thebathymetric gradient decreases down the conti-nental rise to the deep-sea bottom, the last andmost extensive area. Depths of 3000e6000 mare found over 74% of the ocean basins with1% deeper. Perhaps the most characteristic

DEEP OCEAN 17

aspect of the deep-sea bottom is the variety of itstopography. Before any significant deep oceansoundings were available, the sea bottom wasregarded as uniformly smooth. When detailedsounding started in connection with cablelaying, it became clear that this was not thecase and there was a swing to regarding thesea bottom as predominantly rugged. Neither

view is exclusively correct, for we now knowthat there are mountains, valleys, and plainson the ocean bottom, just as on land. With theadvent of satellite altimetry for mapping oceantopography, we now have an excellent globalview of the distribution of all of these features(e.g., Figure 2.1; Smith & Sandwell, 1997), andcan relate many of the features to plate tectonic

FIGURE 2.7 Turbidity current evidence south of Newfoundland resulting from an earthquake in 1929. Source: FromHeezen, Ericson, and Ewing (1954).


processes (Section 2.2) and sedimentation sour-ces and processes.

2.8. SILLS, STRAITS, ANDPASSAGES

Sills, straits, and passages connect separateocean regions. A sill is a ridge, above the averagebottom level in a region, which separates onebasin from another or, in the case of a fjord(Section 5.1), separates the landward basin fromthe sea outside. The sill depth is the depth fromthe sea surface to the deepest part of the ridge;that is, the maximum depth at which directflow across the sill is possible. An oceanic sill islike a topographic saddle with the sill depthanalogous to the saddlepoint. In the deep ocean,sills connect deep basins. The sill depth controlsthe density of waters that can flow over the ridge.

Straits, passages, and channels are horizontalconstrictions. It ismost common to refer to a straitwhen considering landforms, such as the Strait ofGibraltar that connects the Mediterranean Seaand the Atlantic Ocean, or the Bering Strait thatconnects the Bering Sea and the Arctic Ocean.Passages and channels can also refer to subma-rine topography, such as in fracture zones thatconnect deep basins. Straits and sills can occurtogether, as in both of these examples. Theminimum width of the strait and the maximumdepth of the sill can hydraulically control theflow passing through the constriction.

2.9. METHODS FOR MAPPINGBOTTOM TOPOGRAPHY

Our present knowledge of the shape of theocean floor results from an accumulation ofsounding measurements (most of which havebeen made within the last century) and, morerecently, using the gravity field measured bysatellites (Smith & Sandwell, 1997). The earlymeasurements were made by lowering a weight

on a measured line until the weight touchedbottom, as discussed in Chapter S1, SectionS1.1 located on the textbook Web site http://booksite.academicpress.com/DPO/; “S” denotessupplemental material. This method was slow;in deep water it was uncertain because itwas difficult to tell when the weight touchedthe bottom and if the line was vertical.

Since 1920 most depth measurements havebeen made with echo sounders, which measurethe time taken for a pulse of sound to travelfrom the ship to the bottom and reflect back tothe ship. One half of this time is multiplied bythe average speed of sound in the seawaterunder the ship to give the depth. With present-day equipment, the time can be measured veryaccurately and the main uncertainty over a flatbottom is in the value used for the speed ofsound. This varies with water temperature andsalinity (see Section 3.7), and if these are notmeasured at the time of sounding an averagevalue must be used. Research and military shipsare generally outfitted with echo sounders androutinely report their bathymetric data to datacenters that compile the information for bathy-metric mapping. The bathymetry along theresearch ship track in Figure 2.5b was measuredusing this acoustic method.

The modern extension of these single echosounders is a multi-beam array, in which manysounders are mounted along the bottom of theship; these provide two-dimensional “swath”mapping of the seafloor beneath the ship.

Great detail has been added to our knowledgeof the seafloor topography by satellite measure-ments. These satellites measure the earth’sgravity field, which depends on the local massof material. These measurements allow mappingof many hitherto unknown features, such as frac-ture zones and seamounts in regions remotefrom intensive echo sounder measurements,and provide much more information about thesefeatures even where they had been mapped(Smith & Sandwell, 1997). Echo soundermeasurements are still needed to verify the

METHODS FOR MAPPING BOTTOM TOPOGRAPHY 19

gravity-based measurements since the materialon the ocean bottom is not uniform; for instance,a bottom shaped by extensive sediment covermight not be detected from the gravity field.The bathymetry shown in Figure 2.5c, as wellas in the global and basin maps of Figures 2.1and Figures 2.8e2.12, is a blended product ofall available shipboard measurements and satel-lite-based measurements.

2.10. BOTTOM MATERIAL

On the continental shelf and slope most of thebottom material comes directly from the land,either brought down by rivers or blown by thewind. The material of the deep-sea bottom isoften more fine-grained than that on the shelfor slope. Much of it is pelagic in character, thatis, it has been formed in the open ocean. Thetwo major deep ocean sediments are “red”clay and the biogenic “oozes.” The former hasless than 30% biogenic material and is mainlymineral in content. It consists of fine materialfrom the land (which may have traveled greatdistances through the air before finally settlinginto the ocean), volcanic material, and meteoricremains. The oozes are over 30% biogenic andoriginate from the remains of living organisms(plankton). The calcareous oozes have a highpercentage of calcium carbonate from the shellsof animal plankton, while the siliceous oozeshave a high proportion of silica from the shellsof silica-secreting planktonic plants and ani-mals. The siliceous oozes are found mainly inthe Southern Ocean and in the equatorialPacific. The relative distribution of calcareousand siliceous oozes is clearly related to thenutrient content of the surface waters, withcalcareous oozes common in low nutrientregions and siliceous oozes in high nutrientregions.

Except when turbidity currents deposit theirloads on the ocean bed, the average rate of depo-sition of the sediments is from 0.1 to 10 mm per

1000 years, and a large amount of informationon the past history of the oceans is stored inthem. Samples of bottom material are obtainedwith a “corer,” which is a 2e30 m long steelpipe that is lowered vertically and forced topenetrate into the sediments by the heavyweight at its upper end. The “core” of sedimentretained in the pipe may represent materialdeposited from 1000 to 10 million years permeter of length. Sometimes the material islayered, indicating stages of sedimentation ofdifferent materials. In some places, layersof volcanic ash are related to historical recordsof eruptions; in others, organisms characteristicof cold or warm waters are found in differentlayers and suggest changes in temperature ofthe overlying water during the period repre-sented by the core. In some places gradationsfrom coarse to fine sediments in the upwarddirection suggest the occurrence of turbiditycurrents bringing material to the region withthe coarser material settling out first and thefiner later.

Large sediment depositions from rivers createa sloping, smooth ocean bottom for thousands ofmiles from the mouths of the rivers. This iscalled a deep-sea sediment fan. The largest, theBengal Fan, is in the northeastern Indian Oceanand is created by the outflow from many riversincluding the Ganges and Brahmaputra. Otherexamples of fans are at the mouths of the Yang-tze, Amazon, and Columbia Rivers.

Physical oceanographers use sediments tohelp trace movement of the water at the oceanfloor. Some photographs of the deep-sea bottomshow ripples similar to a sand beach after thetide has gone out. Such ripples are only foundon the beach where the water speed is high,such as in the backwash from waves. Weconclude from the ripples on the deep-oceanbottom that currents of similar speed occurthere. This discovery helped to dispel the earliernotion that all deep-sea currents are very slow.

Sediments can affect the properties ofseawater in contact with them; for instance,


silicate and carbonate are dissolved from sedi-ments into the overlying seawater. Organiccarbon, mainly from fecal pellets, is biologicallydecomposed (remineralized) into inorganiccarbon dioxide in the sediments with oxygenconsumed in the process. The carbon dioxide-rich, oxygen-poor pore waters in the sedimentsare released back into the seawater, affectingits composition. Organic nitrogen and phos-phorus are also remineralized in the sediments,providing an important source of inorganicnutrients for seawater. In regions where alloxygen is consumed, methane forms frombacterial action. This methane is often storedin solid form called a methane hydrate. Vastquantities (about 1019 g) of methane hydratehave accumulated in marine sediments overthe earth’s history. They can spontaneouslyturn from solid to gaseous form, causing subma-rine landslides and releasing methane into thewater, affecting its chemistry.

2.11. OCEAN BASINS

The Pacific Ocean (Figure 2.8) is the world’slargest ocean basin. To the north there is a phys-ical boundary broken only by the Bering Strait,which is quite shallow (about 50 m) and 82 kmwide. There is a small net northward flowfrom the Pacific to the Arctic through this strait.At the equator, the Pacific is very wide so thattropical phenomena that propagate east-westtake much longer to cross the Pacific than acrossthe other oceans. The Pacific is rimmed in thewest and north with trenches and ridges. Thisarea, because of the associated volcanoes, iscalled the “ring of fire.” The EPR, a major topo-graphic feature of the tropical and South Pacific,is a spreading ridge that separates the deepwaters of the southeast from the rest of thePacific; it is part of the global mid-ocean ridge(Section 2.2). Fracture zones allow somecommunication of deep waters across the ridge.Where the major eastward current of the

Southern Ocean, the ACC (Chapter 13), encoun-ters the ridge, the current is deflected.

The Pacific has more islands than any otherocean. Most of them are located in the westerntropical regions. The Hawaiian Islands and theirextension northwestward into the EmperorSeamounts were created bymotion of the Pacificoceanic plate across the hotspot that is nowlocated just east of the big island of Hawaii.

The Pacific Ocean has numerous marginalseas, mostly along its western side. In the NorthPacific these are the Bering, Okhotsk, Japan,Yellow, East China, and South China Seas inthe west and the Gulf of California in the east.In the South Pacific the marginal seas are theCoral and Tasman Seas and many smallerdistinct regions that are named, such as theSolomon Sea (not shown). In the southern SouthPacific is the Ross Sea, which contributes to thebottom waters of the world ocean.

The Atlantic Ocean has an “S” shape(Figure 2.9). The MAR, a spreading ridgedown the center of the ocean, dominates itstopography. Deep trenches are found just eastof the Lesser Antilles in the eastern Caribbeanand east of the South Sandwich Islands. TheAtlantic is open both at the north and the southconnecting to the Arctic and Southern Oceans.The northern North Atlantic is one of the twosources of the world’s deep water (Chapter 9).One of the Atlantic’s marginal seas, the Mediter-ranean, is evaporative and contributes highsalinity, warm water to the mid-depth ocean.At the southern boundary, the Weddell Sea isa major formation site for the bottom waterfound in the oceans (Chapter 13). Othermarginal seas connecting to the Atlantic arethe Norwegian, Greenland, and Iceland Seas(sometimes known collectively as the NordicSeas), the North Sea, the Baltic Sea, the BlackSea, and the Caribbean. The Irminger Sea isthe region southeast of Greenland, the LabradorSea is the region between Labrador and Green-land, and the Sargasso Sea is the open oceanregion surrounding Bermuda. Fresh outflow

OCEAN BASINS 21

from large rivers such as the Amazon, Congo,and Orinoco Rivers form marked low-salinitytongues at the sea surface.

The Indian Ocean (Figure 2.10) is closed off byland just north of the tropics. The topography ofthe Indian Ocean is very rough because of theridges left behind as the Indian plate movednorthward into the Asian continent creatingthe Himalayas. The Central Indian Ridge andSouthwest Indian Ridge are two of the slowestspreading ridges on Earth. (As discussed previ-ously, seafloor roughness from abyssal hills andfracture zones is highest at slower spreadingrates, which is necessary in understanding thespatial distribution of deep mixing in the globalocean.) The only trench is the Sunda Trench

where the Indian plate subducts beneath Indo-nesia. The eastern boundary of the Indian Oceanis porous and connected to the Pacific Oceanthrough the Indonesian archipelago. Marginalseas for the Indian Ocean include the AndamanSea, the Red Sea, and the Persian Gulf. The openocean region west of India is called the ArabianSea and the region east of India is called the Bayof Bengal.

The differential heating of land and ocean inthe tropics results in the creation of the monsoonweather system. Monsoons occur in many pla-ces, but the most dramatic and best describedmonsoon is in the northern Indian Ocean(Chapter 11). From October to May the North-east Monsoon sends cool, dry winds from the

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FIGURE 2.8 Map ofthe Pacific Ocean. Etopo2bathymetry data fromNOAA NGDC (2008).


continental land masses in the northeast overthe Indian subcontinent to the ocean. Startingin June and lasting until September, the systemshifts to the southwest monsoon, which bringswarm, wet rains from the western tropicalocean to the Indian subcontinent. While these

monsoon conditions are best known in India,they also dominate the climate in the westerntropical and South Pacific.

Most of the rivers that drain southward fromthe Himalayas d including the Ganges, Brah-maputra, and Irawaddy d flow out into the

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FIGURE 2.9 Map of the Atlantic Ocean. Etopo2 bathymetry data from NOAA NGDC (2008).

OCEAN BASINS 23

Bay of Bengal, east of India rather than into theArabian Sea, west of India. This causes thesurface water of the Bay of Bengal to be quitefresh. The enormous amount of silt carried bythese rivers from the eroding Himalayan Moun-tains into the Bay of Bengal creates the subsur-face geological feature, the Bengal Fan, whichslopes smoothly downward for thousands of

kilometers. West of India, the Arabian Sea, RedSea, and Persian Gulf are very salty due to thedry climate and subsequent high evaporation.Similar to the Mediterranean, the saline RedSea water is sufficiently dense to sink to mid-depth in the Indian Ocean and affects waterproperties over a large part of the Arabian Seaand western Indian Ocean.

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FIGURE 2.10 Map of the Indian Ocean. Etopo2 bathymetry data from NOAA NGDC (2008).


The Arctic Ocean (Figure 2.11) is sometimesnot regarded as an ocean, but rather as a medi-terranean sea connected to the Atlantic Ocean.It is characterized by very broad continentalshelves surrounding a deeper region, which issplit down the center by the Lomonosov Ridge.These shelf areas around the Arctic are calledthe Beaufort, Chukchi, East Siberian, Laptev,

Kara, and Barents Seas. The Arctic is connectedto the North Pacific through the shallow BeringStrait. It is connected to the Nordic Seas (Norwe-gian and Greenland) through passages on eitherside of Svalbard, including Fram Strait betweenSvalbard and Greenland. The Nordic Seas areseparated from the Atlantic Ocean by thesubmarine ridge between Greenland, Iceland,

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Mackenzie R.

FIGURE 2.11 The Arctic Ocean. Etopo2 bathymetry data from NOAA NGDC (2008).

OCEAN BASINS 25

and the UK, with a maximum sill depth of about620 m in the Denmark Strait, between Green-land and Iceland. Dense water formed in theNordic Seas spills into the Atlantic over thisridge. The central area of the Arctic Ocean isperennially covered with sea ice.

The Southern Ocean (Figure 2.12) is notgeographically distinct from the Atlantic,

Indian, and Pacific Oceans, but is often consid-ered separately since it is the only regionoutside the Arctic where there is a path for east-ward flow all the way around the globe. Thisoccurs at the latitude of Drake Passage betweenSouth America and Antarctica and allowsthe three major oceans to be connected.The absence of a meridional (north-south)

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Macquarie Ridge

FIGURE 2.12 The Southern Ocean around Antarctica. Etopo2 bathymetry data from NOAA NGDC (2008).


boundary in Drake Passage changes thedynamics of the flow at these latitudescompletely in comparison with the rest of theocean, which has meridional boundaries.Drake Passage also serves to constrict thewidth of the flow of the ACC, which mustpass in its entirety through the passage. TheSouth Sandwich Islands and trench east ofDrake Passage partially block the open circum-polar flow. Another major constriction is thebroad Pacific-Antarctic rise, which is theseafloor spreading ridge between the Pacificand Antarctic plates. This fast-spreading ridgehas few deep fracture zones, so the ACC must

deflect northward before finding the only twodeep channels, the Udintsev and Eltanin Frac-ture Zones.

The ocean around Antarctica includespermanent ice shelves as well as seasonal seaice (Figures 13.11 and 13.19). Unlike the Arcticthere is no perennial long-term pack ice; exceptfor some limited ice shelves and all of the first-year ice melts and forms each year. The densestbottom waters of the world ocean are formed inthe Southern Ocean, primarily in the Weddelland Ross Seas as well as in other areas distrib-uted along the Antarctic coast between theRoss Sea and Prydz Bay.

OCEAN BASINS 27

introduction to descriptive physical oceanography -...

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