ATER is the most abundant substance on the Earth’s surface, and its properties have a profound effect on the ocean’s chemical, physical, and biological makeup. To understand the behavior of the ocean we need to understand sea water. In this unit we start with the oceanic processes dominated by incoming solar radiation.

Most people live within a few hundred miles of the sea, and all of us live within the ocean’s influence. The ocean provides us with recreation and food, receives our wastes, and serves as our global highway. It is the source from which the atmosphere draws its water. It stores, and later releases, much of the solar energy that powers the winds and causes our weather. By contrasting the large daily temperature changes of an inland desert with the limited temperature range in a coastal climate, we see how deeply the ocean influences our lives.

The unique properties of water are attributed to its molecular structure. Two-hydrogen atoms are combined with one oxygen atom. (NOTE: Both hydrogen atoms share the single electron of the oxygen atom, and are attached at a 105 degree angle to the oxygen atom.) The oxygen side carries a negative charge, the hydrogen side is positively charged. This positive and negative characteristic causes it to be dipolar. The dipolar nature of water gives it solvent properties, surface tension, and high specific heat value.

• Solvent properties–the ability to dissolve other substances.
• Surface tension–the ability to support heavier objects.
• Specific heat–the ability to absorb great amounts of heat.

Lowering the temperature causes the molecular structure to shrink and align in six-sided rings. These rings form the crystal structure of ice. (NOTE: This open network of molecules causes the expansion of water when it changes state from liquid to frozen.) Density increases with depth due to the combined effects of decreasing temperature, increasing salinity and pressure.

We know the major source of heat for the oceans is the Sun. However, the amount of heat energy received, absorbed, and retained by a particular portion of the ocean depends on many things.

The Earth’s curvature causes solar heat energy to be unevenly dispersed. In lower latitudes, the Suns rays generally strike the ocean more perpendicular to its surface. This decreases the area they cover, and increase their ability to heat the ocean. At higher latitudes, sunshine strikes the ocean less perpendicular to its surface. This increases the area it covers, and decreases its ability to heat the ocean.

The percent of the Sun’s electromagnetic energy (all wavelengths) reflected by the surface of an object back into space is known as albedo. The albedo of water varies with the angle of incidence. The albedo for water is 100 percent when the Sun is very low on the horizon, and decreases to 2 percent when the Sun is directly overhead. This decrease is not linear. After the Sun is more than 25 percent above the horizon, the albedo is less than 10 percent. The angle of incidence varies with the latitude of the water surface, the season of the year, and the time of day. Generally, the albedo of water is quite low. The albedo of the ocean varies in some parts of the world due to prolonged periods of cloud cover, large areas of sea ice, and intense atmospheric pollution.

Water is essentially opaque to all electromagnetic radiation except visible light and extremely low frequency (ELF) radio waves. Most energy coming to the Earth from the Sun is in the visible to infrared (IR) portion of the electromagnetic spectrum. Most heat energy is in the IR portion, and water is nearly opaque to it. Most IR energy is absorbed by water before reaching a depth of one meter, so most of the Sun’s heat energy is absorbed at or near the ocean surface. Visible light and ELF radio waves do not play a significant role in heating the ocean’s surface.

The daily input of solar energy is quickly mixed through near-surface waters by wind-induced wave action. Since the heat gained is spread throughout the upper layer of water, it is not easily lost after sunset. This mixing also prevents the ocean surface temperature from rising significantly during the day. On land, the heat can not penetrate below the surface, so it is quickly lost after sunset.

The amount of heat needed to raise the temperature of one gram of a substance 1C° is called specific heat. Several times as much heat is required to raise the temperature of water 1°C than is necessary to raise the temperature of an equal mass of granite. You may notice this difference on a summer day at the beach. The ground becomes hot during the day and cools quickly at night, but water temperatures change very little over a 24-hour period. The specific heat of sea water decreases slightly as salinity increases. The ocean’s capacity to store heat is much greater than that of land due to the high specific heat of water. Therefore, ocean water can absorb and release large amounts of heat and yet change temperature very little. This, in part, accounts for the land having a much greater temperature range than the sea, which results in monsoons and the familiar land and sea breezes of Tropical and temperate regions.

Heat is lost from the ocean surface night and day, and in all seasons, by three processes.

The absorbed short-wave solar heat energy increases the water temperature. The warm water then radiates long-wave radiation back out to space. This is the same process commonly seen on land. About 40 percent of the heat received by oceans is lost to outer space by this process.

The ocean surface is usually about 1°C warmer than the overlying atmosphere. Since the ocean and atmosphere are in contact at the air-ocean interface, heat energy passes from the warmer ocean to the colder atmosphere. About 10 percent of ocean cooling is through this process.

Since the ocean is normally warmer than the overlying atmosphere, there is a natural tendency for water to evaporate. The evaporated water (water vapor) absorbs heat energy from the atmosphere and ocean below to obtain the needed energy to change from liquid to gas. About 50 percent of ocean heat loss is due to evaporation of water. When the water vapor condenses and falls as precipitation, the atmosphere gains heat from the process of condensation. This results in a net temperature gain for the atmosphere because only a portion of the evaporative cooling was atmospheric.

The temperature balance established in the ocean between heat gain and heat loss mechanisms is known as heat budget. For the entire world’s ocean, heat gained equals heat lost. Therefore the ocean’s temperature remains relatively constant. Temperatures at specific places around the world depend on whether there is a net heat gain or loss at that specific location. Large ocean currents and prevailing wind patterns transport, or redistribute, large amounts of energy around the world’s oceans.

Due to the high specific heat of water, ocean temperatures do not vary nearly as much as land mass temperatures at the same latitude. Diurnal temperature variations in the open ocean average 0.2° to 0.3° C. Absolute temperature variations are much smaller than those normally found on land.

The Oceanographic Thermal Equator is a direct reflection of the ocean’s ability to balance temperature. It is located between 5 and 10 degrees North latitude, near, but not coincident with, the thermal equator for the following reasons. More water exists in the Southern Hemisphere (81 percent is covered by water) than in the Northern Hemisphere (61 percent is covered by water). The high specific heat of water allows the Southern Hemisphere to absorb a great deal of heat without a corresponding large temperature increase. The large ice covered continent of Antarctica acts as a cooling mechanism for the entire Southern Hemisphere. The mass transport of water with varying temperatures is accomplished by currents. If there were no surface currents, the sea-surface temperatures would parallel latitude lines.

Just as the air of one region of Earth can differ in its makeup from that of another region, so can sea water. For example, the water around Antarctica differs from that of the mid-latitudes and the Tropics, and water found at the ocean surface differs from that found at or near the bottom. The differences found in sea water are related to sea water properties. It is the sea water properties that are used to classify water masses. In this lesson, we will cover various sea water properties, the three-layer ocean model as it relates to the property of temperature, the water masses of the world’s oceans and land and sea ice.

The properties of sea water have long been a concern to the oceanographer. Whether studying the physical ocean (topography, currents, tides, etc.) or the biological ocean (animal, plant and fish life) they have always been concerned about the properties of sea water and the result of changes to those properties. Since nearly three-quarters of our Earth is covered by water, it’s only logical that changes in the ocean will eventually affect our weather.

Temperature, pressure, and salinity are the three most important properties of sea water, and they determine the other physical properties associated with sea water. This differs from pure water, where only pressure and temperature determine the physical properties. Wave motion and the presence of small suspended particles in sea water are also important variables that affect the properties of sea water. Wave motion causes a change in the processes of chemical diffusion, heat conduction, and transfer of momentum from one layer to another. The suspended particles increase the scattering of radiation, thereby absorbing more radiation than a similar layer (thickness) of pure water. The variables of wave motion and suspended particles, although important, cannot be measured.

Besides temperature, pressure, and salinity, other common physical properties of sea water are water color, transparency, ice and sound velocity. Some of the lesser known properties include specific heat, compressibility, osmotic pressure, eddy viscosity, electrical conductivity, radioactivity, and surface tension. Many of the lesser known properties can only be determined using complex mathematical calculation and formulation that incorporates data on one or mole of the common physical properties, especially temperature, pressure, and/or salinity.

Temperature is the most important property of sea water. The ocean, like the atmosphere, is heated by the Sun’s incoming radiation. In all latitudes the ice-free portions of oceans receive a surplus of radiation. Some of this heat is given up to the atmosphere, and some of it is retained. Because the sea retains a portion of this heat, the sea-surface temperature (SST) is normally higher than the air temperature. However this is true only when average conditions are considered. Whether the sea surface is warmer or colder than the air above it at any particular moment depends on the locality, the season of the year, the character of the atmospheric circulation and the character of the ocean currents.

The temperature of the ocean ranges from about –2°C to 30°C. Ocean water that is nearly surrounded by land may have higher temperatures, but the open sea, where the water is free to move about, hardly ever heats above 30°C. Here, the ocean currents distribute the heat and tend to equalize the temperature. Deep and bottom water temperatures are always low, varying between 4°C and 1°C.

SSTs change from day to night just like those of the atmosphere, but to a much lesser degree. The diurnal variation of SST in the open ocean is on the average only 0.2°C to 0.3°C. The greatest diurnal variation takes place in the Tropics, with lesser variation at higher latitudes. The range of diurnal variation depends on the amount of cloudiness and the direction and speed of the wind.

The smallest seasonal temperature variation of surface water occurs in equatorial and polar regions. The largest seasonal temperature variation occurs in the mid-latitudes (±2° Celsius from the monthly mean). In areas where warm and cold currents meet, surface temperatures may differ by 4°C. The absolute maximum is about 32°C in the Red Sea and Persian Gulf in summer with an absolute minimum of about –2°C in the polar region during the winter.

The annual variation of SST in any region depends on the variation of incoming radiation, the character of the ocean currents, and the character of the atmospheric circulation. The annual range of surface temperatures is much greater over the oceans of the Northern Hemisphere than those of the Southern Hemisphere. This wider range of temperatures appears to be associated with the character of the prevailing winds, particularly the cold winds blowing from the continents. On the other hand, the annual range of ocean temperatures in the Southern Hemisphere is most definitely related to the range of incoming solar radiation, because of the absence of large land masses south of 45°S. Here, the prevailing winds travel almost entirely over water. This causes a greater degree of consistency in the annual SST patterns and a much smaller annual temperature range compared to the northern Hemisphere.

The temperatures near the Equator experience a semiannual variation. This corresponds to the twice yearly passage of the Sun’s most direct rays across the Equator.

The annual variation of temperature in subsurface layers depends on several additional factors–namely, the variation in the amount of heat that is directly absorbed at different depths, the effect of heat conduction, the variation in currents related to lateral displacement, and the effect of vertical motion. Diurnal temperature variations in subsurface layers are largely unknown. What we do know is that they are extremely small.

As we move away from the upper layer of the ocean surface, where the heat from the Sun is stored, and toward the ocean floor, the temperature of the ocean drops (fig. 1–1). This gives us the vertical temperature distribution. There are three layers in the ocean regarding temperature: The mixed layer is the layer closest to and inclusive of the surface; The main thermocline; and the deep-water layer.

Due to more heat input into lower latitudes than higher latitudes, it can be assumed that sea water temperatures decrease poleward. Sea water temperatures are warmer in each hemisphere’s summer.

Pressure is the second most important property. Pressure beneath the sea surface is measured in decibars. The pressure exerted by one meter of sea water very nearly equals one decibar (1/10 of a bar) or 100,000 dynes per square centimeter.

In the ocean for each 32.8 feet (10 meters) increase in depth there is an equivalent increase of one atmosphere of pressure. This may also be expressed as a one bar increase for every ten-meter increase in depth. The decibar unit is a convenient measurement because it very nearly equals the corresponding depth in meters.

The farther one descends in the sea, the greater the pressure, and since pressure in the ocean is essentially a function of depth, the numerical value of pressure in decibars approximates the ocean depth in meters. Therefore, pressure ranges from zero at the surface to over 10,000 decibars in the deepest parts of the oceans. The pressure is created by the weight of the sea water above. The weight per unit volume of sea water, in turn, varies with the temperature and salinity. In a column of water of constant depth, the pressure increases as the temperature of the sea decreases, or the salinity increases, or both.

Salinity is the third most important property. The term salinity is related often to the amount of salt in the water. In oceanography, salinity is defined as ‘‘the total amount of dissolved solids in sea water." Salinity is measured in parts per thousand by weight, and is symbolized by ‰. The measurement gives us the grams of dissolved material per kilogram of sea water.

The salinity values of ocean water range between 33‰ and 37‰ with an average of 35‰. In the open ocean, surface salinity is decreased by precipitation, increased by evaporation, and changed by the vertical mixing and inflow of adjacent water. Near shore, salinity is generally reduced by river discharge and freshwater runoff from land. In the colder waters that freeze and thaw, salinity generally increases during periods of ice formation and decreases during periods of ice melt.

Salinity is at its minimum in the fall. This corresponds with Tropical storm season. The vast amount of precipitation associated with these Tropical storms as they move across the ocean surface decreases salinity.

Latitudinally, surface salinity varies in a similar manner in all oceans. Maximum salinity values occur between 20° and 23°N and S, whereas minimum salinity values occur near the Equator and toward higher latitudes. The controlling factor in average surface salinity distribution is the latitudinal differences in evaporation and precipitation. Exceptions to this statement do occur, and local variations should be expected, especially near the mouth of the larger river systems and in the Atlantic coastal water of the United States, Labrador, Spain, and Scandinavia. The best known region of strong horizontal salinity gradients is the Grand Banks region, where warm, saline Gulf Stream water mixes with the colder, less saline water of the Labrador

Here, water with a salinity value as low as 32‰ may override or lie next to water having a salinity value greater than 36‰. A similar situation prevails in the Pacific Ocean, where the Kuroshio and Oyashio currents mix.

At latitudes poleward of 40°N and S, where precipitation generally exceeds evaporation, salinity values tend to increase with depth. Usually during summer, these positive salinity gradients are accompanied by strong negative temperature gradients and result in very stable water, especially in the coastal regions. These strong, shallow salinity (and temperature) gradients persist through the summer.

In the open oceans, salinity is normally considered a constant. The arctic and coastal regions experience an input of large quantities of fresh water in the spring when ice melt or river runoff reduces the salinity of the surface layer. This becomes a transition layer, creating a "salinity front". High salinity levels occur where evaporation exceeds precipitation such as the Eastern Mediterranean (39‰) and the Red Sea (41‰).

The density of sea water depends on salinity, temperature, and pressure. At constant temperature and pressure, density varies with salinity. A temperature of 32°F and an atmospheric pressure of 1013.2 millibars are considered standard for density determination. At other temperatures and pressures the effects of thermal expansion and compressibility are used to determine density. The density at a particular pressure affects the buoyancy of various objects, notably submarines. Density is defined as mass per unit volume, and is expressed in grams per cubic centimeter.

Density and temperature have an indirect (inverse) relationship. As temperature increases density decreases. Likewise, if temperature decreases density increases. Density has a direct relationship with both pressure and salinity. As pressure and/or salinity increases density increases.

The greatest changes in density of sea water occur at the surface. Here, density is decreased by precipitation, runoff from land, melting of ice, or heating. When the surface water becomes less dense, it tends to float on top of the denser water below. There is little tendency for the water to mix; therefore, the condition is one of stability. The density of surface water is increased by evaporation, the formation of sea ice, and cooling. If the surface water becomes denser than the water below, it sinks to a level having the same density. Here, it tends to spread out to form a layer, or to increase the thickness of the layer of which it has become a part. As the denser water sinks, the less dense water rises, and a convective circulation is established. The circulation continues until the density becomes uniform from the surface to a depth at which a greater density occurs. If the surface water becomes sufficiently dense, it sinks all the way to the bottom. If this occurs in an area where horizontal flow is unobstructed, the water that has descended spreads to other regions, creating a dense bottom layer. Since the greatest increase in density occurs in polar regions, where the air is cold and great quantities of ice form, the cold, dense polar water sinks to the bottom and then spreads to lower latitudes. This process has continued for such a long time that the entire ocean floor is covered with this dense polar water. This explains the layer of cold water at great depths in the ocean.

Compressibility is the ability of water to be compacted under pressure. Sea water is not as compressible as pure water, due to salinity and increased specific heat generated during the compression process. Sea water is highly elastic making it a good medium for sound. This value changes slightly with changes in temperature or salinity. The effect of compression is to force the molecules of the substance closer together, causing the substance to become denser. Even though the compressibility of sea water is low, the total effect is considerable because of the amount of water involved. If it were zero, sea level would be about 90 feet higher than it is now.

Viscosity is resistance to flow. Sea water is slightly more viscous than freshwater, and the level of resistance is controlled by its temperature and salinity. Viscosity increases when salinity increases or the water temperature decreases. However, the effect of decreasing temperature is greater than that of increasing salinity. The resistance rate is not uniform; it increases as the temperature decreases. Because of the effect of temperature on viscosity, an incompressible object might sink at a faster rate in warm surface water than in colder subsurface water. For most compressible objects, viscosity effects may be more than offset by the compressibility of the object. In reality this is a very simple explanation to a complex problem, since the actual relationships existing in the ocean are considerably more complicated than portrayed here.

Liquids expand and contract when temperature changes take place; some more than others. Sea water has a higher coefficient of expansion than that of freshwater. Within the sea, the coefficient of thermal expansion is affected by salinity, temperature, and pressure. It is greater in high salinity water; greater in warm water than in cold (under similar salinity conditions); and it increases with increasing depth under constant temperature and salinity conditions. Of course, constancy is not a trademark of any of these properties; they are all quite variable. In turn, the thermal expansion that takes place in the sea varies and is difficult to assess.

A major role of thermal expansion is in the formation of ice. Pure water is densest at 4°C. Thermal expansion takes place when water warms above 4°C, but it also expands when it cools below 4°C. When expansion takes place, the volume is increased, which in turn decreases the density. When water cools below 4°C, it expands slightly, and as it freezes, it expands much more. If water failed to expand during the freezing process, the density of ice would be such that it would sink to the bottom on forming. In the cold of winter, freshwater lakes would eventually become solid blocks of ice. Come summer, only the upper few feet of ice would melt, leaving the remaining ice beneath the melted water.

The property that enables a substance to change its length, volume or shape in direct response to a force, and recover its original form upon removal of that force. Directly proportional to sound speed and inversely proportional to compressibility.

A convenient method of visualizing the sea is to divide it into layers in much the same way that we do the atmosphere. Using bathythermograph information (temperature versus depth profiles), the oceans display a basic three-layered structure: the mixed layer, main thermocline, and deep-water layer. The latitudinal distribution of these layers is shown in figure 1–2, while the typical thermal structure is shown in figure 1–3. Both figures are representative of winter.

The mixed layer is the upper layer of the three-layered ocean model. It is a layer of fairly constant warm temperatures that, in middle latitudes, extends from the surface to a maximum depth of about 450 meters, or 1,500 feet. This layer gets its name from the mixing processes that cause its fairly constant warm temperatures. The two mixing processes are classified as mechanical and convective.

This mixing process is caused by wave action, surface storms, etc. The wave action stirs up the water. Warmer surface water is driven downward, where it mixes with colder subsurface water. Eventually, a layer of water with a fairly constant temperature is produced. This process is more important in summer than in winter,

because surface waters are much warmer and less dense than subsurface waters, thereby producing a stable water column. The mechanical mixing process is more rapid and irregular than the convective mixing process.

This process occurs as a result of changes in water stability. When surface waters become denser than subsurface waters, an unstable condition exists. Such a condition can occur when there is an increase in surface salinity owing to evaporation or the formation of ice, or by a decrease in the surface water temperature. A temperature decrease of 0.01°C or a salinity increase of 0.01‰ is sufficient to initiate the convective mixing process. In the former case, for example, a cold polar or arctic air mass moving over warm water cools the surface water before it can cool the subsurface water. As the surface waters cool and become colder than the subsurface waters, they become denser and sink. As the colder surface water sinks, the warmer and less dense subsurface water rises to the surface to replace it. This process continues until the water is thoroughly mixed, the density difference eliminated, and the water column stabilized.

Even though winds and the resultant wave action are generally stronger during winter, convective mixing, caused by the colder winter air temperatures, produces a deeper mixed layer than can be attained by mechanical mixing. It is for this reason that convective mixing is considered the more important of the two, and the predominant process in winter.

The convection process is strongest in northern waters where vertical temperature and salinity gradients are not extreme and surface waters undergo a high degree of cooling. Convective mixing attributed to salinity changes is most noticeable in the Mediterranean and Red seas, where evaporation far exceeds precipitation.

We have looked at both processes individually; however, the two processes can and often do take place simultaneously. When this occurs, the mixed layer normally attains a greater depth than would be attained by either process individually.

This is the part of the ocean where temperature decreases rapidly with depth. Mixing no longer affects the ocean layers, and especially the solar heating of the surface.

The main thermocline is the central layer of the ocean generally between 1,000 and 3,000 feet. It is found at the base of the mixed layer and is marked by a rapid decrease of water temperature with depth. The portion of the ocean where temperature commonly decreases rapidly with a small increase in depth; no significant changes occur from season to season.

At high latitudes there is no marked change in water temperature with the seasons, while in the mid-latitudes, a seasonal thermocline develops with the approach of summer. This seasonal thermocline comes about from the gradual warming of the surface water during spring. The warming takes place in the upper few-hundred feet of the surface, and results in the seasonal thermocline becoming superimposed on the main thermocline. Figure 1–4 illustrates the development of the seasonal thermocline in the mid-latitudes.

The mid-latitude summer thermocline is more pronounced than the thermocline of spring or autumn. Bathythermograph traces of the summer thermocline show that it affects a much broader range of depth than at any other time of year. In our illustration, the seasonal thermocline is roughly 35 meters thick (90 to 125 meters deep). NOTE: That the winter temperature profile shows no seasonal thermocline. The mixed layer extends to a depth of more than 160 meters. Come spring, the surface water is warmed and a seasonal thermocline develops between 35 and 60 meters. As summer takes hold, the water warms to 25 °C and the mixed layer extends to a depth of approximately 90 meters. The thermocline now exists between 90 and 125 meters. In summer, the seasonal thermocline is deeper and covers a broader range of depth than at any other season of the year.

With the approach of autumn, the mixed layer continues to drive the thermocline deeper, but the water within the mixed layer is cooler than it was in summer. Just as in the spring, the cooler water in the mixed layer decreases the range of depths covered by the thermocline. In low latitudes, small seasonal temperature changes make it difficult to distinguish between the seasonal and the permanent thermoclines.

A small decrease in vertical temperature, typically 1° to 2° C, observed above the permanent thermocline in the mixed layer; short-term phenomenon usually associated with heating and/or wind mixing. The stronger the wind/wave action the deeper the transient thermocline–the tradeoff is a reduction in magnitude since the absorbed heat is spread over a larger area in the mixing process. The afternoon effect (caused by diurnal heating) results in a transient thermocline.

The deep-water layer is the bottom layer of water, which in the middle latitudes exists below 1,200 meters. This layer is characterized by fairly constant cold temperatures, generally less than 4°C.

To better understand the basic vertical temperature distribution, look again at figure 1–3. At high latitudes in winter, the water is cold from top to bottom. The vertical temperature profile is essentially isothermal (no change in temperature with depth). In middle latitudes, the structure is like that illustrated in figure 1–4. In low latitudes, the mixed layer extends to a depth of about 300 feet. Here, the main thermocline is encountered and the temperature drops about 1°C more than it does in the mid-latitudes. This sharper drop is due to the higher surface temperature in the lower latitudes. The thermocline extends to 2,100 feet, where the deep layer is encountered.

The idea of visualizing water masses as we do air masses is possible because both are based on the physical properties that go into their makeup. The properties of temperature and salinity are used to classify both water types and water masses.

A water type has a single value of salinity and a single value of temperature associated with it, while a water mass takes into account a range of temperatures and salinity’s. For example, Red Sea water is a water type characterized by a temperature of 9°C and a salinity of 35.5‰. On the other hand, North Atlantic Central Water (a water mass) is characterized by a range of temperatures (4°C to 17°C) and salinity (35.1‰ to 36.2‰). A water mass may be considered to be made up of a combination of two or more water types.

The vast majority of water masses are formed at the surface of the sea in middle and high latitudes. Cold, highly dense surface water sinks until it reaches a level having the same constant density. Here, it spreads out horizontally. The manner in which it spreads out depends on its density in relation to the density of the surrounding water. This is true of nearly all water masses, except those of low latitudes–in particular, the equatorial water masses of the Indian and Pacific Oceans. These water masses are formed by the mixing of subsurface waters.

Water masses can be modified through mixing with other masses being advected or from atmospheric intervention (seasonal or diurnal). Normally these changes are negligible. NOTE: A typhoon near Guam in the mid-eighties became stationary for three days. At the end of the three day period it dissipated leaving behind a cold pool of surface water. This atmospheric intervention caused substantial upwelling to drastically alter the water mass in that area.

In low and middle latitudes the vertical arrangement of water is such that we can distinguish a surface layer, upper water (central and equatorial), intermediate water, deep water, and in some localities, bottom water. In high latitudes, the layered structure all but disappears because the surface water is similar to the water at or near the bottom.

The surface layer is not classified as a water mass or water type, because its properties vary widely from one area to another, depending on current variations, evaporation, precipitation, and various seasonal changes, especially in the middle latitudes. In low

and middle latitudes it is found above central and/or equatorial water to depths of 100 to 200 meters. The surface layer is separated from deeper water by a transition layer (the main thermocline).

Beneath the surface layer, we come across the water types and water masses. Like air masses, the water types and water masses have source regions in which they form. Figure 1–5 is provided as a reference for the source regions of various water types and water masses.

Central water is normally found in relatively low latitudes although its source region is around the subtropical convergences (between the 35^th and 40^th parallel in each hemisphere). Convergence’s are regions in the ocean where surface waters are brought together by the currents. In the western North Atlantic Ocean, a region of subtropical convergence exists where the Gulf Stream meets the colder, denser Labrador Current. Convergences are marked by rapidly rising SSTs.

Central water is not usually discernible at the surface and is generally relatively shallow. Its greatest thickness is observed along its western boundaries. In the western North Atlantic around the Sargasso Sea, the thickness may reach 900 meters.

Variations in heating and cooling, evaporation and precipitation, ocean circulation patterns, and mixing processes all contribute to the salinity values of central water being either quite similar or considerably different. For example, central water of the South Atlantic Ocean, the Indian Ocean, and the western South Pacific Ocean all have similar salinity values, while the salinity values of North Atlantic central water are considerably higher than the central water of the North Pacific Ocean.

You will note as you look at figure 1–5 that the central water of the North and South Atlantic oceans is not separated by equatorial water like the central water of the North and South Pacific oceans. Instead, the central water of the North and South Atlantic come together and mix, forming a region of transition consisting of intermediate properties.

Equatorial water is found in the Pacific and in the Indian Ocean. In the Pacific it is thought to originate on the southern side of the Equator. There are two reasons for this: Its properties are similar to those of the water masses of the South Pacific, and its salinity values are higher than those of the water masses found in the North Pacific Ocean.

Equatorial water is also found in the northern part of the Indian Ocean. Here, its higher salinity’s are probably due to its mixing with the waters of the Red Sea. However, this conclusion has not been proved.

Equatorial water, like central water, is not discernible at the surface, because the temperature and salinity values used to isolate it cannot be clearly ascertained in the upper 100 to 200 meters.

Intermediate water is found below central water in all oceans. Intermediate water includes Antarctic intermediate water, Arctic intermediate water, Mediterranean water, and Red Sea water.

Antarctic intermediate water encircles the Antarctic continent and is the most widespread of all the intermediate water masses. It forms near the Antarctic convergence, where it sinks. As it sinks, it flows north and mixes with the water masses that lie immediately above and below it.

In the Atlantic, the absence of equatorial water allows Antarctic intermediate water to flow across the Equator and reach roughly 20° to 35° N latitude. In the South Pacific and Indian oceans, where equatorial water does exist, Antarctic intermediate water fails to reach the Equator. It spreads north to about 10°S latitude.

One of the characteristics of Antarctic intermediate water is its low salinity (34.1‰, to 34.6‰). In comparison to the water around it, it displays the lowest salinity values.

Arctic intermediate water and Sub-Arctic water are similar; however, in the North Atlantic Ocean, Arctic intermediate water forms only in small quantities, and in a relatively small area east of the Grand Banks of Newfoundland.

In the North Pacific, Arctic intermediate water forms during winter at the convergence formed by the Oyashio current and the Kuroshio Extension. It exists between latitude 20° and 43°N, except off the west coast of North America. Here, Sub-Arctic water extends to lower latitudes, and the northern boundary of the intermediate water is pushed much farther south.

This water mass is formed by the interaction of dense Mediterranean Sea water with the waters of the adjacent North Atlantic Ocean. The denser Mediterranean water flows out through the Strait of Gibraltar and sinks to a depth of about 1,000 meters, where it mixes with the water at this depth.

This water type is found over large parts of the equatorial and western regions of the Indian Ocean. Large quantities of warm, highly saline water from the Red Sea flow into the Indian Ocean, where its mixes with Antarctic intermediate water to form the Red Sea water mass. The spreading of Red Sea water is not as well defined as Mediterranean water.

This water mass is thought to form through a combination of mixing and vertical circulation in the region between the subtropical and Antarctic convergences. Here, large quantities of Antarctic intermediate water and Antarctic bottom water mix with North Atlantic deep water to form Antarctic circumpolar water.

The physical properties of this water mass are quite conservative, and as its name implies, it extends completely around the Antarctic continent and the South Pole. Because Antarctic circumpolar water forms in the deeper waters of the Antarctic Ocean, it is often referred to as Sub-Antarctic water.

Sub-Arctic water is much like Antarctic circumpolar or Sub-Antarctic water; however, there are differences. The differences are attributed to the land and sea distribution in the two hemispheres. In the Southern Hemisphere, the Antarctic convergence extends around the continent of Antarctica, but in the Northern Hemisphere, the Arctic convergence is found only in the western portions of oceans. However, even in these areas the convergence is not always well defined.

In the North Atlantic Ocean, Sub-Arctic water covers a relatively small area, and it possesses a higher salinity than surrounding waters. On the other hand, the Sub-Arctic water of the North Pacific is much more extensive, and its salinity values are lower than surrounding waters.

In the deep-ocean basins below intermediate water, high density deep and bottom water exists. These water masses form in both hemispheres. In the Southern Hemisphere, Antarctic bottom water forms near the Antarctic continent, while in the Northern Hemisphere, Arctic deep and bottom water forms in northwestern Labrador Basin and in a small area off the southeast coast of Greenland. These water masses form at the surface, sink, and spread out to fill the deep-ocean basins. Deep and bottom waters are detectable in areas far removed from their source regions.

Roughly three percent of the world’s water surfaces are ice covered. In the polar seas, submarines transit beneath the ice-covered water, while surface ships operate in and around the ice. From a safety standpoint, all ships must be aware of where ice lies in relation to their position.

The two main types of ice found in the seas and oceans are sea ice, which accounts for 95 percent of the total coverage, and glacier ice. The Naval Polar Oceanography Center at Suitland, Maryland, keeps the US Navy advised of the development, movement, and the equatorward limit of sea ice, and the location and movement of icebergs. This is done through the issuance of global sea-ice analyses and forecasts.

We know that pure water freezes at 0°C, but the freezing point of sea water varies depending on salinity (fig. 1–6). Under average salinity conditions (35‰), sea water begins to freeze at –1.9°C. However, before surface waters will freeze, the entire water column must cool to its temperature of maximum density. Therefore, shallow waters (of low salinity–less than 24.7‰) freeze more rapidly than

deeper water. As you can see, the freezing of sea water is governed primarily by temperature, salinity, and depth; however, this ice formation can be retarded by winds, currents, and tides.

In the open sea, the first sign that the sea surface is freezing is an oily opaque appearance of water. This appearance is caused by the formation of minute ice needles (spicules) and thin plates of ice (frazil crystals). As the formation process continues, the surface attains a thick soupy consistency termed grease ice. Next, depending on the wind, waves, and salinity, an elastic or brittle crust forms. The elastic crust (nil’s) has a matte appearance, while the brittle crust (ice rind) is shiny. As the crust thickens, the wind and sea cause the ice to break up into rounded masses known as pancake ice. With continued freezing, the pancake ice forms into a continuous sheet.

Sea ice forms by the freezing of sea water at temperatures near –2° C. Sea ice is made up of crystals of pure ice separating small cells of brine. As the ice further cools salt solids in the brine crystallize out leaving pure ice and increasing the salinity of the surrounding water. If the ice temperatures’ raise to near 0°C the brine and encompassing water will melt first and become hummocked leaving behind ice that can be used as potable water.

Sea ice properties differ greatly from those of fresh water ice. Properties depend on:

• The amount of enclosed brine.
• Number of air bubbles left in the ice if all or part of the brine has become hummocked (trickled down).

Density of pure ice at 0° C is 0.9168. The density of sea ice may be more or less than that amount depending on the brine and/or air bubble amount.

The formation of sea ice usually begins with the onset of autumn, and the first ice usually appears in the mouths of rivers that empty into shallow seas, such as that off northern Siberia. During the increasingly longer and colder nights of autumn, ice

forms along the shorelines (fast ice) and becomes a semipermanent feature that widens and spreads. When islands are close together, as in the Siberian Sea, fast ice blankets the sea surface, and bridges the waters between all land areas.

On the average in the Northern Hemisphere, sea ice is at a minimum in September, while at a maximum in March. In the Southern Hemisphere these times are nearly opposite; minimum in March and maximum in September.

There are only certain areas of the Earth where sea ice forms that we will discuss.

The high latitudes encompass the polar regions, the polar (or frigid) zones, defined by the Arctic Circle (66.3°N) to 90°N, and the Antarctic Circle (66.33°S) to 90°S. These boundaries approximate the 50°F isotherm during the warmest months.

Six countries border the Arctic: the United States, Canada, Soviet Union, Norway, Finland and Greenland. Nine ocean areas (approximately 4,600,000 square nautical miles) are contained within the Arctic: the Arctic Ocean, Beaufort Sea, Chukchi Sea, Kara Sea, Barents Sea, Greenland Sea, Laptev Sea, East Siberian Sea, and the Norwegian Sea. Permanent sea ice covers most of the Arctic Ocean north of 85°N.

This area includes all the arctic countries plus Iceland, Sweden, Ireland, United Kingdom, Poland, and Germany. It includes all the Arctic Ocean areas plus the Bering Sea, Sea of Okhotsk, North Sea, Labrador Sea, extreme northern Atlantic and Pacific Oceans, and the Hudson Bay area. Seasonal sea ice (pack ice) extends south over the region to approximately 40°N in the winter. Pack ice concentrations depend on latitude, wind, and cloud cover.

This area includes the Antarctic continent and all the peripheral islands. It includes the extreme southern Indian Ocean, Atlantic and Pacific Oceans, Bellingshausen Sea, Ross Sea, Weddell Sea, Amundsen Sea, and the Davis Sea. Permanent ice covers the Antarctic continent and extends out along the continental shelf from 15 to 40 nautical miles.

It includes the Antarctic Palmer Peninsula, the tip of Chile and Argentina, the Antarctic Ocean areas and the Tasman Sea.

The amount of snow and/or ice cover in any region of the Arctic determines the radiation budget of that region.

Solar absorption over snow-free land may be relatively high and as a result, temperatures in the summer may rise to 40° to 50°F. Over the ice pack solar absorption is much less, keeping temperatures at 32°F year round. At the ice edges, temperature and moisture gradients is often quite large. Greater than normal precipitation, cloudiness, fog and altered cyclone tracks occur at the sea-ice interface. Leads and polynyas have a similar effect, they provide a significant source of heat and moisture. Open water produces a moderating effect on local temperatures, while snow covered land or ice produces extremely cold temperatures. Actual snowfall amounts are small due to the lack of water vapor at low temperatures. Arctic ocean averages five inches annually, while the ocean margin areas average approximately ten inches annually. Blowing snow is prevalent in the high latitudes and creates the illusion of more snowfall than is actually occurring. NOTE: The average depth of snow cover in March to April range from 8 to 20 inches over the frozen oceans to 16 to 28 inches over the subarctic regions. The effect of the variation in snow cover is reflected in the fact that there is a very short transitional period between winter and summer. (Spring and fall last only a few weeks.)

Permanent ice fields exist over Greenland in the Arctic and over the continent of Antarctica in the Southern Hemisphere. Permanent ice may be sea ice or land-fast ice created from eons of snow accumulation and glaciation. Greenland has an elevation over 11,000 feet and the vast ice plateau of Antarctica is over 14,000 feet AGL. Icebergs "calve" off the land ice and begin their travel into the adjoining ocean following the wind and currents. Icebergs can travel up to 12 miles in a day. The predominant currents’ control the speed and heading the icebergs travel. Wind and the coriolis effect also influences the iceberg’s movement. The portion of the iceberg above the water is the "freeboard" and it acts like a sail, the larger the freeboard the more the wind affects the iceberg’s speed.

Sea ice is subdivided into young ice, first-year ice, and old ice.

This ice forms in one year or less, and its thickness ranges from 10 to 30 centimeters (4 to 12 inches). It is further classified as gray ice and gray-white ice.

This ice is a reasonably unbroken level of ice of not more than one winter’s growth that starts as young ice. Its thickness is from 30 centimeters to 2 meters (1 foot to 6 ½ feet). First-year ice may be subdivided into thin first-year ice, medium first-year ice, and thick first-year ice. The latter is more than 4-feet thick.

Old ice is extremely heavy sea ice that has survived at least one summer’s melt. It occurs primarily in the arctic and Antarctic polar packs as a vast mass of converging, crushing, and dividing ice floes of various ages, sizes, shapes, and thicknesses that drift around the Arctic Basin and Antarctica. Old ice may be subdivided into second-year ice and multi-year ice.

Generally, sea ice is categorized into seven sizes. Refer to figure 1–7 for relative sizes and a comparison to other more common features.

The terms most frequently used to describe the topography, or configuration of the ice surface, are related to the degree of surface roughness. Figure 1–8 illustrates the types of topography.

This type of topography occurs when ice cakes override one another. Rafting occurs when wind forces ice cakes or ice floes together. It is associated with young and first-year ice.

Ridged ice is much rougher than rafted ice and occurs with first-year ice. Wind and weather can eventually smooth the surface of the ridges.

Pack ice, sea ice covering more than half of the visible sea surface, usually drifts to the right of the true wind in the Northern Hemisphere (left in the Southern Hemisphere). Observations show that the actual drift is about 30 degrees from that of the wind direction, or very nearly parallel to the isobars on a weather product. The drift more closely follows the wind in winter than in summer. In summer, the tides play a bigger role in the movement of the ice.

A close estimate of the speed of drifting pack ice is possible using the wind speed. On the average, the drift of ice in the Northern Hemisphere ranges from 1.4 percent of the wind speed in April, to 2.4 percent of the wind speed in September. Although wind is the primary driving force, the presence or absence of open water in the direction of the drift greatly influences the speed of drift. Ice-free water in the direction of the drift, no matter how distant, permits the pack ice to drift freely in that direction. Ice-clogged water, on the other hand, slows the ice’s forward movement.

Naval operations in and around fields of sea ice can be hazardous. The movement of massive floes of ice can cut off ships from open water, and worse yet, the ice may close in around a ship, leaving it stranded in a sea of ice. Therefore, open water in ice-covered seas becomes very important.

There are a great variety of water features associated with sea ice. Some of the most common features are as follows:

1. Fracture. Any break through the ice.
1. Lead. A long, narrow break or passage through the ice; a navigable fracture. A lead may be open or refrozen.
1. Puddle. A depression in sea ice usually filled with melted water caused by warm weather.
1. Thaw holes. Holes in the ice that are caused by the melting associated with warm weather.
1. Polynya. Any sizable area of sea water enclosed by sea ice. Put another way, simply a large hole in the ice.

Ice of land origin composed initially of large accumulations of compacted snow that reach the sea as coastal glaciers and ice shelves. The leading edges of these glaciers break off (calve) and fall into the sea. This ice then drifts to sea as icebergs.

Since 86 percent of the world’s glaciers occur in Antarctica, most icebergs originate around that continent. Most of the remainder of the world’s glaciers are located in Greenland. Greenland is the main source of icebergs in the Northern Hemisphere (about 90 percent). Nearly 70 percent of Greenland’s icebergs originate along the western coast near 68°N.

Icebergs are pinnacled (cone-shaped) or tabular (flat-topped and straight-sided) as shown in figure 1–9. The structure, and to some extent the appearance, depends on the ice that produces the berg. Pinnacled or irregular-shaped bergs come from glaciers that plow across uneven ground on their way to the tidewater, while the tabular bergs

come from ice shelves that thrust directly out to sea. Pinnacled and irregular bergs are most prevalent in the Northern Hemisphere, while the tabular bergs are more prevalent in the Southern Hemisphere.

Icebergs originating in Greenland average 70 meters in height and 280 to 450 meters in length when first formed. The largest ones may exceed 120 meters in height and several miles in length. The tabular bergs of Antarctica average 30 to 40 meters in height, but their horizontal dimensions greatly surpass those bergs of the Northern Hemisphere. For example, one iceberg observed near Scott Island in 1956 measured 60 miles by 208 miles.

The portion of an iceberg that is visible above the water depends on the type of the berg and the density differences between the sea water and the ice. The type of berg (pinnacle or tabular) determines the height of the ice above the water. In the case of the tabular berg, the depth below the surface is about seven times the height above the water line. In the case of the pinnacle or irregular berg, the depth below the surface averages about five times that above the water line.

Regarding density, sea water with a temperature of –1°C and a salinity of 35‰ produces a density condition that allows for nearly 90 percent of the ice to be submerged.

Irregular icebergs often have rams (protrusions of ice beneath the surface). These rams can be a great hazard to vessels that might pass close to this type of bergs.

While the general direction of the drift of icebergs over a long time is known, it may not be possible to predict the drift of an individual berg at a given place and time, for bergs lying close together have been observed to move in different directions. The reason for this is that icebergs move under the influence of the prevailing current at the iceberg’s submerged depth. This subsurface current often opposes the existing wind and sea or surface drift.

Like icebergs, bergy bits and growlers originate from glaciers and form when icebergs and other masses of land ice disintegrate. A bergy bit is a medium sized fragment of

glacier ice and is about the size of a small cottage. A growler is a small fragment of ice about the size of a grand piano. It is usually of glacial origin, and generally greenish in color.

The Earth’s topography (mountains, valleys, etc.) has a definite and important effect on the elements and characteristics of its surrounding atmosphere. This relationship exists between the ocean floor and the oceans. The irregular terrain of the ocean floor affects the movement of ocean water, temperature gradients in areas of channeling, and in the area of naval operations, submarine and antisubmarine warfare (ASW) tactics. Many relief features and bottom types are used by submariners to conceal their submarines and lessen their chances of being detected by surface sonars. The surface fleet must also be aware of the relief features and bottom types to assess the effectiveness of their search sonars.

If we were able to walk from the land above sea level to the deepest depths beneath the sea, our walk would be mostly downhill. On leaving the shore, our walk would take us through five major bottom provinces; the continental shelf, the continental slope and rise, the ocean basin, and the mid-ocean ridges. Figure 1–10 shows the different undersea geophysical features.

The continental shelf is the first province we come across on leaving land. The average width of the shelf is approximately 40 miles, but in some places, there is no shelf (for example, along the west coast of South America). The widest shelf is found along the glaciated coast of Siberia, where it extends out roughly 800 miles. Continental shelves comprise about 7.5 percent of the total ocean bottom.

The shelf has a very gradual slope. It declines at an average rate of 2 fathoms per mile, and at its seaward limit, the water above the shelf is usually 6 to 100 fathoms deep (1 fathom = 6 feet). Although the average slope of the shelf is gradual, terraces, ridges, hills, depressions, and deep canyons are found within its boundaries.

The shelf region is a transition zone between freshwater runoff from land and the more saline water of the sea; consequently, it is an area of great mixing of water with generally unstable water conditions. Currents normally run parallel to the shore in this region.

Continental slope

At the seaward edge of the continental shelf the slope becomes much steeper. This region is known as the shelf break. The drop off is rapid. On the average, the slant ratio is roughly 20 times greater than that of the continental shelf. The ratio is generally much greater off mountainous coasts than off wide, well-drained plains. On bottom contour charts (fig. 1–11), the bottom contours are tightly packed, thereby reflecting the much steeper gradient.

The continental slope resembles a steep cliff that has been eroded by heavy rains. Its most striking features are the submarine canyons and deep cuts or scars that are prevalent along the slope face. These canyons are thought to have been formed (or cut out) by turbidity currents, which are dense, sediment-laden currents that flow along the ocean floor. Some of these canyons are equal in size to the Grand Canyon.

At the seaward end of these canyons, large amounts of sediment are deposited and spread out in a fan-like manner to form the continental rise. Refer to figure 1–10 again.

The continental rise is found seaward of the continental slope, in approximately 500 fathoms of water. It is made up of thick sediment deposits that cover irregular relief features. These deposits slope gently seaward forming the abyssal plains of the deep ocean basins. At the seaward edge of the continental rise, the water depth is about 1,500 fathoms.

The ocean basins account for 76 percent of the ocean floor, and their depths range from 1,500 to 3,000 fathoms. They have a very slight average incline of no more than 1:90 miles. For every 90 miles seaward the bottom slopes no more than one mile. However, superimposed on this very flat plain are many rugged relief features, such as seamounts, guyots, atolls, volcanic islands, sills, and trenches.

Seamounts are submerged, isolated, pinnacled mountains rising 3,000 feet or more above the sea floor. Generally volcanic in nature they most often have smaller bases in proportion to their height. Due to the small base, the concentration of mass causes them to sink (over a long time) leaving an atoll. Seamounts are more common in the Pacific. They may exist alone, in chains, or in groups but not in a line.

Guyots are submerged, isolated, flat-topped mountains that rise 3,000 feet or more above the sea floor. Guyots are more abundant in the Pacific than the Atlantic.

Atolls are seamounts or guyots that have broken the surface, and coral deposits have built up around the rim. The coral forms a reef around a shallow body of water–a lagoon. Atolls may form on a subsiding continental shelf or around the margin of a sinking volcanic island in open ocean areas.

These islands occur individually and in groups (island arcs). They are formed by volcanic eruptions. About 10,000 volcanoes dot the ocean floor, and they are especially abundant in the western Pacific basin. The Hawaiian Islands are probably the best known example of volcanic islands. In the North Atlantic Ocean, one of the most recent volcanic island (Surtsey) was formed south of Iceland along the Mid-Atlantic Ridge.

Sills are elevated parts of the ocean floor that partially separate ocean basins. A sill restricts the movement of bottom water masses and results in their partial, and in some cases nearly total, isolation.

Trenches are long, narrow, and relatively steep-sided depressions. They comprise the deepest portions of the oceans. The trenches of the Pacific Ocean stretch for as long as 2,500 miles (Peru-Chile Trench), are more numerous than in any other ocean, and have by far the greatest depths in the oceans. For example, the Mariana Trench is 35,800 feet deep; the Tonga Trench, 35,430 feet deep; and the Mindanao Trench, 34,428 feet deep. Trenches are normally found on the seaward side of island arcs, while relatively shallow seas exist on the continental side.

On leaving the abyssal plains, we come to the last of the oceanic provinces, the mid-ocean ridges. The Mid-Atlantic Ridge is the most conspicuous of all ridges. It extends from Iceland southward across the Equator to about 55°S, forming an eastern and western basin in the Atlantic. The Mid-Atlantic Ridge rises from a depth of 2,500 fathoms and is continuous at depths of less than 1,500 fathoms over the greater part of its length. In several places, this ridge rises above sea level to form islands such as the Azores and Ascension.

Note to Student: Consider all choices carefully, select the best answer to each question, and circle the corresponding letter. When you have completed all unit review exercises, transfer your answers to ECI Form 34, Field Scoring Answer Sheet.