Chapter 13: The Seasons of the Sea

We now turn our attention to changes in sunlight incident on Earth’s surface that occur as a result of Earth’s orbit around the Sun, what we know as the seasons. These seasonal changes in solar radiation have a profound effect on nature and human activities. The seasonal cycle represents one of the most dynamic and important processes on Earth. The term “season” originated from French and Latin words meaning “to sow or plant.” People in ancient times depended on the planting and harvesting of crops for survival, as they do in modern times. Seasons marked the times when certain farming activities had to be completed. Time off from school during summer—the so-called summer break—arose from the need for young people to help with the planting and harvesting of crops. While fewer modern-day students work on farms than in times past, the tradition of summer break continues to this day. (Aren’t we lucky?) 

In the modern age, winter, spring, summer, and fall have special significance in our lives for the clothes we wear, the activities we do, the foods we eat, and the holidays we celebrate. We look forward to the changing seasons. Winter brings snowboarding, spring means Coachella, summer ushers in days at the beach, and fall allows you to dress up like a ghoul. Each new season brings an opportunity to look forward to something, a chance to say, “Next season, I’m going to . . . ”

The ocean has its seasons, too, its own Valentine’s Days, Earth Days, Fourth of Julys, and Halloweens. As the seasons bring changes in the intensity and duration of sunlight, the ocean’s surface cools or warms. Sea ice forms or melts. Plant and animal life diminish or flourish. Like the changing colors of foliage on land, so do the seasons bring changing colors of the ocean, too. In Southern California, the seasonal cycle brings gray whales in winter. It draws beach-nesting grunion fish to our shores in spring, puts forth dazzling tide pool life in summer, and lights the nighttime surf in a blue glow in fall. 

In many ways, the seasons of the sea illustrate the most spectacular story of the ocean and its interconnections with life and our planet. Your understanding of the reasons for the seasons provides not only a framework for how the ocean works as a system but also an awareness of its many moods. Each season drives profound shifts in the physics, chemistry, biology, and ultimately the geology of our planet. By the end of this chapter, you’ll have a better appreciation for how and why.

13.1 Sunrise, Sunset

Two very important and obvious-if-you-think-about-them phenomena mark the seasons. First, the time of the sunrise and the time of the sunset change daily. Unless you arrived on Planet Earth only recently, the nights of your entire life have gotten longer during certain times of the year and shorter during other times. Now, some people will say that the days get shorter in the winter and longer in the summer. But it’s really more accurate to say that daylength, defined as the lighted portion of the day, changes over the seasons. In actuality daylength increases from the first day of winter until the first day of summer. From the first day of summer to the first day of winter, the daylength hours diminish.

Second, the point on the eastern horizon where the Sun comes up and the point on the western horizon where it goes down change every day. (You knew that, right?) These points actually have a name: the sunrise point, the location on the eastern horizon where the Sun rises, and the sunset point, the location on the western horizon where the Sun sets. The reason for the change in the sunrise and sunset points has to do with the path of the Sun as it crosses the sky. As it moves from the sunrise point to the sunset point, the Sun traces an arc—the solar path. And the solar path changes daily. Now, the solar path might not be something that you’ve ever taken notice of, but you’ve likely noticed changes in the sunrise and sunset points. Have you ever been nearly blinded by the rising or setting sun while driving along an east–west highway? That happens because the sunrise or sunset points fall in a direct line with the road you’re on. You’ve fallen victim to the seasonally changing solar path.

13.2 Why Do Seasons Occur?

Day and night length and sunrise and sunset points change because of Earth’s position as it orbits the Sun. To understand Earth’s orbit, we need to review a few astronomy terms and concepts. But before we start, let’s clear up one very common misconception about the seasons.

Earth’s tilt (explained below) causes the seasons. As you learned in Chapter 12, the Sun heats Earth’s surface (with shortwave radiation). Earth’s surface, in turn, heats our atmosphere (with longwave radiation). Because Earth is tilted with respect to the incoming solar radiation, some places receive more intense sunlight than others. Some times of the year, a given location will receive direct sunlight. Other times of the year, it will receive less direct sunlight. The result is intense heating some times of the year and less intense heating effect other times of the year. Those changes in heating we know as the seasons. 

Much of this chapter is devoted to exploring the details of the seasonal cycle so don’t worry if it’s not clear to you just yet. But if the tilt model of Earth’s season doesn’t satisfy you, ask yourself this: Why does the Northern Hemisphere enjoy summer in June while folks in the Southern Hemisphere experience winter in June. Let’s find out why. 

13.2.1 Earth’s Rotation

As I’m sure you know by now, Earth turns like a disco ball in outer space, in what is called Earth’s rotation. The imaginary line around which it rotates is called Earth’s axis of rotation, or simply Earth’s axis. Grab a spherical piece of fruit—an orange works nicely—and carefully jab a pencil through the middle of it. (Styrofoam balls work for this too, but not everyone keeps Styrofoam balls in their house.) Hold the top end of the pencil–orange physical model of Earth so that the pencil is vertical. Then twirl the orange. Voilà! You have created a model of how Earth rotates on its axis. Now add this detail: Look down on your model from above the North Pole (where the North Pole is oriented up) and rotate it counterclockwise. Your model now correctly represents the direction of rotation of Earth around its axis. The Sun rises in the east because the Earth rotates towards the east. That means the Sun rises first on the east coast of the US and even earlier in the UK.  

13.2.2 Earth’s Tilt

Next, using a marker, draw an equator at the halfway point between the top and bottom of the orange (i.e., midway between the North and South Poles). Tilt the pencil to a vertical position (straight up and down). The equator you just drew should be horizontal. Now—and this is the most critical part of your demonstration—try a 45-degree tilt (symbolized as 45°), halfway between vertical and horizontal. (Review degrees and angles on the internet if this terminology is unfamiliar to you.) Then try a 22.5° angle, halfway between vertical and 45°. This is close to the angle of Earth’s axial tilt, the angle between vertical and the axis of rotation of a planet. (Astronomers officially call this obliquity, but that word is harder to remember, so we’re sticking with tilt.) Earth’s axial tilt is roughly 23.5° in modern times. (To be precise, it’s 23.43681°, but most textbooks and websites round it off to 23.5, apparently because it’s easier to remember, but what could be easier to remember than 2-3-4?) 

13.2.3 Earth’s Orbit

Now we need to see what happens when Earth orbits the Sun, what’s called Earth’s orbital revolution or Earth’s orbit. For this, you will need a model of the Sun—a lamp, a grapefruit, a chair in the middle of the room, anything that your body can orbit. Start with Earth held so that its axis of rotation is at a 23.5° angle with the North Pole pointed away from your sun. Hold it firm in that position while you walk in a circle around your imaginary sun. When you have returned to your starting position, note the position of your Earth. It should be the same: at a 23.5° angle pointed away from your sun. As you orbited your sun holding your tilted Earth, you traced an orbital plane, the horizontal plane in which Earth’s orbit occurs. The equator also can be considered as a plane, called the equatorial plane. Viewed in this way, the equatorial plane makes a 23.5° angle with the orbital plane. Alternatively (just looking at it a different way), Earth’s tilt represents the angle between Earth’s axis of rotation and a vertical line drawn perpendicular to the orbital plane. Earth’s tilt is maintained throughout its orbit around the Sun, 365.25 days a year.

Now, if you want to be really accurate, your simulation of Earth’s orbit would take a slightly elliptical shape rather than a perfect circle. As it turns out, the distance between Earth and the Sun varies during Earth’s annual orbit around the Sun. You may recall this distance as the astronomical unit, or AU. By definition, the Earth–Sun distance is 1 AU, about 93 million miles. Although the AU is considered a constant, we know the Earth is slightly closer to the Sun in January. At this location, known as perihelion—when a planet is closest to its star—Earth is positioned about 91.4 million miles away from the Sun. In July, at aphelion—when a planet is farthest from its star—Earth is slightly farther away, about 94.5 million miles (e.g., NASA 2001). This should put to rest any notion that the Earth–Sun distance causes the seasons.  The Sun is closer to the Earth in January—when the Northern Hemisphere experiences winter—and farther from the Earth in July during Northern Hemisphere summer.

13.3 How Does Earth’s Tilt Affect Surface Warming?

What does all this mean? As Earth orbits the Sun with its equatorial plane tilted at an angle, the Sun’s rays—which you can imagine as arrows—strike it at an angle. Some parts of Earth heat more at certain times of the year—those that receive more direct sunlight—and some parts heat less—where the Sun’s rays arrive at an angle. When the Northern Hemisphere is tilted toward the Sun and the intensity of sunlight is high, we experience summer. When the Northern Hemisphere is tilted away from the Sun and the intensity of sunlight is low, we experience winter. It’s this phenomenon—the tilt of Earth as it orbits the Sun—that creates the differences in heating that we notice as the seasons. 

You can model this with a flashlight and a piece of paper. First, turn the flashlight on and hold it directly above the piece of paper. Draw a circle around the lighted region of the paper. Now tilt the flashlight so that the light strikes the paper at an angle. Look at the illuminated part of the paper. Has the circle changed shape? (Yes!) It should now look more elliptical. Draw an ellipse around the lighted portion of the paper. Now compare the circle and the ellipse. Which one covers a greater area of the paper? Hopefully, you can see that the ellipse covers a greater area. If you were as tiny as an insect and stood inside the circle, the light would appear brighter than when you stood inside the ellipse. Although the same total amount of light falls on the circle and the ellipse, the light in the ellipse is spread over a greater area. The light per unit area is greater in the circle than in the ellipse. Sunlight striking Earth’s surface at an angle will be less intense than sunlight striking Earth’s surface directly. It’s like splitting a pizza with more people. With more people, each person gets less pizza. 

Earth’s tilt explains the change in daylength and the change in the sunrise and sunset points on the eastern and western horizons, respectively. An observer in the Northern Hemisphere sees the Sun low in the sky when the Northern Hemisphere is pointed away from the Sun. When the Northern Hemisphere is pointed toward the Sun, an observer sees the Sun high in the sky. The Northern Hemisphere observer also sees that the sunrise and sunset points are farther south in the winter and more northerly in the summer. The solar path across the sky in winter is lower than the path across the sky in summer. Put another way, the solar altitude angle, or simply, the noon sun angle—the position of the Sun above the equatorward (southern in Northern Hemisphere) horizon at noon—is lowest in winter and highest in summer. Stick a toothpick in your orange at the latitude of Southern California—about 30°N—to visualize how the sun angle changes as you orbit the Sun. With the sun higher in the sky from winter to summer, daylength is longer. The opposite is true from summer to winter. The seasonal changes in daylength, the sunrise and sunset points, the solar path, and the pattern of warming and cooling are explained simply as a result of Earth’s tilt relative to its plane of orbit around the Sun. 

Now, before we proceed, it’s important that you recognize that the seasons of the Northern Hemisphere are not the same seasons for the Southern Hemisphere. In fact, they are exact opposites. When it’s summer in the Northern Hemisphere, it’s winter in the Southern Hemisphere. When it’s spring here, it’s fall there. Most of South America and all of Australia experience winter during our summer. Their beach days come in January instead of July. When the Southern Hemisphere is pointed toward the Sun, the Northern Hemisphere is pointed away. This causes the seasons of the two hemispheres of Earth to occur in different months of the year. That’s why some folks travel south (or north) during their winter. They can enjoy summer almost year round!

13.3.1 Solstices and Equinoxes

The times of year when the Sun is absolutely at its lowest or highest points in the sky are known as the solstices, meaning “when the sun is still.” These are the days when ancient people perceived that the Sun no longer moved lower or higher. We know these days as the winter solstice, the first day of winter and the shortest day (i.e., daylength is at its minimum) of the year, and the summer solstice, the first day of summer and the longest day (i.e., daylength is at its maximum) of the year. 

What happens to daylength on the day following a solstice? Daylength on the day after the winter solstice is slightly longer, and the day after the summer solstice, it’s slightly shorter. In fact, daylength increases from the date of the winter solstice to the date of the summer solstice. On the other hand, daylength decreases from the date of the summer solstice to the date of the winter solstice. If you can remember the shortest day of the year (first day of winter, the winter solstice) and the longest day of the year (first day of summer, the summer solstice), then you can easily figure out what is happening to daylength at any time of the year. 

We also mark the days when the Sun crosses directly overhead at the equator. These are known as the equinoxes (equi = “equal”; nox = “night”), the times of the year when day and night are of equal length.  The moment when the Sun is directly over the equator on its way into the Northern Hemisphere is called the Northern Hemisphere spring equinox or vernal equinox. Half a year later, as the Sun passes southward on its way into the Southern Hemisphere, the moment when it’s directly over the equator is called the Northern Hemisphere fall equinox or autumnal equinox. Of course, the seasons are reversed for the Southern Hemisphere. On the spring and fall equinoxes, you’ll find the Sun directly over the equator. That means if you stood on the equator at local noon on this day (when the Sun is directly overhead), you would barely see a shadow at your feet. “Look, Mom, no shadow!”

13.3.2 The Midnight Sun

Earth’s axial tilt dramatically affects daylength at the North and South Poles. When the Northern Hemisphere is tilted away from the Sun—during the Northern Hemisphere fall and winter—sunlight can’t reach the North Pole. This region experiences 24 hours of darkness or twilight for six months out of the year. Alternatively, when the Northern Hemisphere is pointed toward the Sun—during the Northern Hemisphere spring and summer—the northern polar region experiences 24 hours of sunlight. People refer to regions poleward of the Arctic and Antarctic Circles as the land of the midnight sun, places where the Sun appears 24 hours a day. The period when this occurs is referred to as the polar day, the six-month-long period of continual (or near-continual) sunlight in these regions. The six-month-long period of darkness in these regions is called the polar night. More than a novelty, polar day and polar night have profound effects on polar environments.

13.3.3 Lines of Latitude Mark the Seasonal Cycle

Astronomers and geographers have long noted the changing positions of the Sun in the sky over the course of a year. These positions have been given special importance by the establishment of names for latitudes of astronomical significance. The place where the Sun is directly overhead on the first day of summer in the Northern Hemisphere, 23.5°N, is called the Tropic of Cancer. Where the Sun is directly overhead on the first day of Northern Hemisphere winter, 23.5°S, we find the Tropic of Capricorn. The latitude above which no sunlight can reach following the fall equinox is known as the Arctic Circle, 66.5°N. At the other end of Earth, the latitude below which no sunlight falls following the Southern Hemisphere fall equinox, we find the Antarctic Circle, 66.5°S. 

We should also note that regions of Earth that receive direct sunlight most times of the year stay warm, and those that receive less sunlight or variable sunlight experience colder temperatures. Because of these differences, we can divide Earth into three simplified climate zones, the average weather for a region of Earth based on the amount of sunlight they receive. The region of Earth bounded by the Tropic of Cancer and the Tropic of Capricorn is known as the tropical zone, or the tropics. This zone experiences the warmest weather on the planet because the Sun is mostly overhead most days of the year. The polar zones—regions north of the Arctic Circle and south of the Antarctic Circle—encompass the coldest regions of the planet. The region between the tropical and polar zones—the middle latitudes of Earth—are known as the temperate zones. Here, where most of the continental United States lives, the climate is moderate—not too hot and not too cold, on average.

The tropical, temperate, and polar climate zones have precise astronomical meaning. Their boundaries refer to extremes in the positions of the Sun (i.e., Tropic of Cancer and Tropic of Capricorn) or its rays (Arctic Circle and Antarctic Circle). The three climate zones also provide a general guide to the kind of climate one might experience. The tropical zone at low latitudes is generally warm, humid, and rather seasonless. The temperate zones at middle latitudes are mild, with the four classic seasons (winter, spring, summer, and fall). The polar zones at high latitudes are cold and dry, with two seasons: the cold season—polar summer and fall—and the colder season—polar winter and spring. Though other configurations of climate regions based on temperature and precipitation patterns (e.g., the Köppen classification) provide more detail and remain the ones preferred by meteorologists and climatologists, the three-climate-zone model serves well for understanding most ocean-related phenomena. The interested reader may consult a good introductory meteorology textbook for fancier models (e.g., Ahrens and Henson 2018).

13.4 The Ocean Layer Cake

Now, to get the most out of this next section, you’ll want to place a multilayered cake in front of you. The concepts here will make more sense if you place a slice of your favorite, delicious, hard-to-resist layer cake in front of you. Even if you don’t like to eat cake, the presence of a slice will help. Avoid eating it for now. We’ll refer to cake throughout the rest of this chapter. And when you finish the chapter, you’ll have a nice reward!

The three climate zones mentioned above—the tropical, temperate, and polar zones—receive varying amounts of sunlight due to Earth’s orbit around the Sun. This means that the parts of the ocean within these zones receive seasonally varying amounts of sunlight. Tropical oceans heat the most; polar oceans heat the least. This differential heating of the planet drives a number of important processes. Here we take a look at how seasonal changes in heating affect the physical structure of the ocean; that is, how heating and cooling cause changes in the temperature structure of the water column. Fear not, if you like swimming in lakes in the summer, this section will be a piece of cake.

13.4.1 Seawater Density (Revisited)

In earlier chapters, you learned that changes in seawater temperature and salinity can cause changes in seawater density, the mass of molecules occupying a given volume of seawater (i.e., mass per unit volume). Here we examine those changes in greater detail.

The exchange of heat across the air–sea interface may bring heating or cooling of the surface of the ocean. The amount of heating or cooling depends on the solar intensity and the temperature of the atmosphere. When sunlight is intense—or the atmosphere is warmer—the surface waters will warm. When sunlight is less intense—or the atmosphere is cooler—the surface waters will cool. Of course, a number of additional factors can affect heating and cooling of the ocean’s surface. Clouds, airborne or waterborne particles, winds, waves, and even the color of the water may affect temperature changes. But in general, sunlight and atmospheric heat are considered the major factors governing ocean heat content (e.g., Lindsey and Dahlman 2020).

Temperature affects seawater density because it affects the spacing between the molecules of water and its dissolved salts. As we heat seawater, its molecules move faster, and the space between the molecules gets larger. As we cool seawater, the molecules move slower, and the molecules get closer together. In the first case—heating—a rise in temperature causes the density of seawater to decrease. We say the seawater becomes less dense. In the second case—cooling—the lowering of temperature causes the seawater to become more dense. Temperature and density have an inverse relationship: As one quantity goes up, the other goes down, and vice versa.

The concentration of salts in seawater—its salinity—also affects its density. As the salinity—the concentration of salts—increases, the density increases. That’s because there are now more molecules packed into the same volume of seawater. If we add freshwater and lower the salinity, we decrease the density. Salinity and density have a positive relationship (or correlation). As one increases or decreases, so does the other.

And we shouldn’t forget that the solid form of water—ice—is less dense than the liquid form of water. That’s because water molecules repel each other and the structure of water changes at temperatures below 39.2°F (4°C). Because ice is less dense than liquid water, it floats. 

Finally, as noted in Chapter 11, freezing of seawater and formation of sea ice cause brine rejection, extrusion of a syrup of salts into the surrounding ocean. The sea ice is fresh, but the brine increases local seawater salinity as it dissolves.

13.4.2 The Concept of Buoyancy

Heating and cooling and increases and decreases in salinity change the density of seawater.  As a result of changes in density, a seawater parcel—an oceanographer’s informal term for an unspecified volume of seawater—may sink or rise. The position of a water parcel in the ocean depends on the buoyant force, the upward force exerted on a fluid or object immersed within it. The buoyant force counteracts the gravitational force—the pull of gravity on the water parcel or objects within it. It’s these two forces that determine buoyancy—the tendency of a fluid or object to rise or sink in a fluid. Buoyancy underlies the different density layers in the ocean—that is, the ocean layer cake.

A brief explanation provides some context for the physics. The buoyant force arises because water pressure—exerted in all directions—increases with depth (as the weight of water above an object increases). The deeper the depth, the greater the buoyant force. The buoyant force also depends on the volume of the parcel or object. The larger the volume, the greater the buoyant force on that object. Objects of equal volumes at the same depth will experience the same buoyant force. However, if their masses differ, the gravitational force on them will differ. That’s because the gravitational force depends on the mass of the object (i.e., the gravitational force equals the mass of the object times the gravitational constant). The mass of a given volume depends on its density: the higher the density, the greater its mass. In short, water parcels with a higher density will experience a greater gravitational force. And this will influence where the water parcel comes to rest in the water column.

Here are the key concepts related to the importance of buoyancy for understanding ocean layering. If the buoyant force is less than the gravitational force, then the fluid or object sinks, which is known as negative buoyancy. If the buoyant force exceeds the gravitational force, then the fluid or object rises, a condition known as positive buoyancy. If the buoyant force and the gravitational force are equal, then the fluid or object remains stationary—it has achieved neutral buoyancy. If you are on a sinking ship, you will want to grab something that will help you stay afloat, something that has positive buoyancy in water. If you want to hide a dead body, then you will want to attach it to something that will sink, something with negative buoyancy. All kidding aside, the principles of buoyancy have important practical applications, such as scuba diving, shipbuilding, and operating a submarine, among others. 

The principles of buoyancy are more obvious with a familiar example, such as a hot air balloon—invented in 1783, making it the oldest form of human flight (Kotar and Gessler 2011). By heating air inside a flexible container—the balloon—it becomes positively buoyant. The air becomes less dense, the balloon gains buoyancy, and it rises. When the hot air cools, either by conduction through the fabric of the balloon or by letting cool air into the balloon, the balloon loses buoyancy and it sinks. But how does a balloon remain stationary at an ideal altitude above the ground? There comes a point in its upward transit where the density (and pressure) of the surrounding air matches the density (and pressure) of the air inside the balloon. At that point, the hot air balloon stops its ascent. The air inside and outside the balloon has the same density, and the balloon achieves neutral buoyancy.

Scuba divers maintain their buoyancy using a combination of weights (negative buoyancy) and a special air-holding vest called a buoyancy compensator, or BC, for short (positive buoyancy). When the diver wishes to descend, they need negative buoyancy, so they release air from their BC. When they wish to maintain a particular depth—to hover over a shipwreck or beautiful reef, for example—they inflate their BC just enough to balance the gravitational force of their weights and so achieve neutral buoyancy. When they want to ascend and return to the surface, they add a little more air to their BC, which causes them to acheive positive buoyancy and rise slowly.

Water parcels also sink and rise in the water column, albeit without weights and BCs. If the water parcel loses heat or gains salts, it may become negatively buoyant and sink. Gaining heat or freshwater, the parcel may become positively buoyant. It rises—unless it’s already at the surface. If the water parcel is neutrally buoyant, it remains in place. Just as they do with hot air balloons and scuba divers, the principles of buoyancy govern the movements of water masses and play a role in the productivity of the ocean.

13.4.3 A Stable or Unstable Water Column

In a given region of the ocean or a lake, the water parcels will naturally arrange themselves according to their density and buoyancy. The least dense and most buoyant water parcels will be at the surface. The most dense and least buoyant water parcels will be at the bottom. We can view these water parcels as density layers—like a layer cake—where the water parcels are arranged according to their density. When the layers are arranged in order of increasing density—least dense on the top and most dense on the bottom—we have a stable water column. If one or more of the layers are out of order—if they are not arranged in order of increasing density from the surface to the bottom—then you have an unstable water column.

Looking at the water column as a stack of water parcels with different densities helps explain many of the phenomena we observe in the ocean or even a lake. Ever swim in a lake during summer and noticed that it’s very cold when you dive deeper? That’s because the lake consists of different layers of water, each with its own density. The top layer is warm, relatively light (i.e., less dense), and positively buoyant. The bottom layer is cold, relatively heavy (i.e., more dense), and negatively buoyant. What seems like a nice, warm summer swim turns into a shockingly cold experience when you take a dive. In a nutshell, the same thing happens in the ocean. 

13.5 Measuring the Ocean’s Layers

Since the time of the Challenger expedition, oceanographers used various types of water sampling bottles for obtaining seawater samples at depth. The Pettersson–Nansen bottle, co-invented by Norwegian oceanographers Otto Pettersson (1848–1941) and Fridtjof Nansen (1861–1930), became commonplace on oceanographic vessels by the early 20th century (e.g., Mill 1900). Typically, a single bottle or a string of bottles is lowered in the open position until the bottles reach the desired depth, whereupon they’re closed, trapping a volume of seawater inside. When the bottles are brought back on board, samples of water are taken for immediate analysis or preserved for future analysis. 

While perfectly fine for many types of chemical analyses, this method introduces problems for measurements of temperature. As the sample is raised, its temperature may change as it exchanges heat with the surrounding seawater. This gives an inaccurate reading of the sample’s actual temperature at depth. The invention of reversing thermometers solved the problem of temperature measurements at depth. These thermometers could be attached to sampling bottles and turned upside down at depth using a weighted messsenger sent down the line. This to cut off the supply of mercury and prevented any further changes in the reading (e.g., Deacon 1971). 

So, too, oceanographers desired a faster means for determining salinity. The salinometer—an electronic instrument for measuring salinity—reduced the need for laborious chemical measurements. Combined, reversing thermometers and salinometers permitted oceanographers to take more samples in more places in the world ocean (e.g., Salinometry 2023). 

Improvements in computing technology and reductions in the size of computers by the early 1970s paved the way for development of a new kind of instrument, the CTD—the conductivity–temperature–depth instrument. Based on an instrument developed by American engineer Neil Brown and colleagues in the early 1960s, the CTD is considered the workhorse of oceanographic research (e.g., Salinometry 2023). There is perhaps no better instrument for obtaining a near-instantaneous measurement of the properties of the water column. If CTDs had public appeal, you can imagine an ad campaign similar to one for a popular credit card: “The CTD—Don’t leave your home port without it.” (See The Drum Team 2016 for the history of the ad from which this popular phrase is borrowed.) 

It didn’t take long for oceanographers to trick out their CTDs. One of the first things they did was attach the instrument package (the CTD) to a frame on which water sampling bottles could be placed. Modern CTDs come equipped with a rosette, a series of water sampling bottles. Most often these are Niskin bottles, invented by American businessman Shale Niskin (1926–1988; Niskin 1962). Arranged in a circle of 12 to 24 bottles and triggered electronically, the Niskin bottles provide a means of taking water samples on the fly. An oceanographer can look at the vertical profile of temperature and salinity and choose a depth to take a water sample simply by raising or lowering the CTD rosette to that depth. Other kinds of electronic sensors can be attached for measurements of submarine light intensity (i.e., radiometers), chlorophyll concentration (i.e., fluorometers), pH (pH probes), dissolved oxygen (oxygen probes), and more. When an oceanographer raises a CTD to go over the side and watches it swinging back and forth as the ship rocks in the waves, they pray that it will come back safely. One cable snap can send a $250,000 fully outfitted CTD rosette to the bottom of the sea—and an oceanographer crying all the way home.

13.6 Visualizing Layers: The XZ Graph

To represent the properties of the water column—such as those measured during a CTD cast—oceanographers have created a special kind of graph. The XZ graph resembles an upside-down XY graph, except that the vertical axis, the z-axis, extends downward from the origin. In an XY graph, the y-axis extends upward from the origin. The vertical z-axis represents water depth, usually in meters. In the top left corner of the graph—the origin, where the x-axis and the z-axis intersect—the depth is set to zero. Thus, the x-axis at zero depth represents some property, X, at the surface of the ocean. Properties such as temperature, salinity, and light often change with depth. An XZ graph illustrates the changes in ocean properties you might experience as you descend into the deep. Multiple x-axes are allowed too. By adding x-axes above the first one to represent other properties (such as salinity, chlorophyll, and others), oceanographers can see how properties vary simultaneously with depth in the water column. (See Hautala for additional details.) 

A useful way to envision the water column and how its properties change with depth is to stare at a ceiling, the higher the better. Imagine that you are looking up at the ocean surface—the ceiling—from the sea bottom—the floor. Then consider how some property, like temperature, might vary from the ceiling (i.e., ocean surface) to the floor (i.e., sea bottom). In your home, hot air rises, so the warmest air accumulates near the ceiling. As you descend from the ceiling, the air becomes cooler. The coldest air can be found at floor level. And so it is with the ocean. Warm surface waters lie atop deeper, cold waters.

13.7 The Seasons and the Thermoclines

We now come to my honest-to-goodness favorite topic in oceanography: the layering of the ocean due to seasonal heating by the Sun. For me, it’s one of the best examples of how the ocean works as a system. The physics affects the biology, which in turn affects the chemistry of the ocean. Physics drives ocean life, which alters the ocean’s chemistry. But that’s getting ahead of our story. In this chapter, we’ll focus mostly on the physics. And we’ll look just at temperate oceans for the time being. But be prepared. We’ll return to these topics later in the book and explore the marvelous symphony of processes conducted by the Sun. 

13.7.1 Winter and the Permanent Thermocline

The different layers that we observe in the ocean result largely from seasonal differences in heat exchange. Salinity plays an important role too—especially in the polar oceans—but we’ll concentrate on heating first. The temperate zone ocean provides the best example of how changes in heat exchange at the surface affect the layering. And the best place to start is winter.

In the winter, the ocean generally loses heat to the atmosphere because the overlying atmosphere is colder. Surface heating is low because daylengths are short and sun angles are low. The loss of heat from the ocean to the atmosphere cools the ocean’s surface. A cold, dense, negatively buoyant layer of surface water forms. Because this cold surface layer rests on top of a warmer, less dense, and more positively buoyant layer, the water column becomes unstable. The surface layer begins to sink.

Sinking of surface waters continues throughout the winter, as long as the top layer gets colder than the layer beneath it. At some point—when the coldest seasonal air temperatures have been reached—the coldest surface layer forms. This layer represents the most negatively buoyant layer formed during the season. As a result, this layer will sink the deepest.

The sinking of surface waters in winter represents an important process by which the upper ocean is mixed. Mixing is just what it sounds like (minus the DJ): the ocean is blended, combined, homogenized, and stirred. Any water parcels with slightly different temperatures (or salinities) may be mixed by the sinking action of denser and thus less buoyant water parcels. This is called buoyancy-driven mixing. It’s important for distributing heat, gases, salts, biologically important nutrients, particles, plankton, larvae, and all sorts of other things in the upper ocean. This wintertime mixing often results in an isothermal water column (iso = “same”; thermal = “temperature”), one in which temperature does not change with depth. This layer of mixed waters in the upper ocean is called—wait for it—the surface mixed layer. The distance from the surface to the bottom of the mixed layer is called the mixed layer depth. The mixed layer depth gives an estimate of the depth of mixing of the water column.

How deep is the mixed layer? Negatively buoyant surface waters sink until they reach a depth where the surrounding water has the same density—they sink until they become neutrally buoyant. In tropical oceans, the surface layer may sink a few tens of meters in winter (e.g., Longhurst 1993). In temperate oceans, the surface layer may sink from 15 to 75 meters on average (e.g., Bathen 1972). In polar environments, where the coldest air temperatures occur, the surface layer may sink thousands of meters, even all the way to the seafloor (e.g., Johnson 2008). The densest waters in the world ocean are formed in polar oceans. These abyssal waters—as they are collectively called—spread along the seafloor from the poles to the equator and beyond. Abyssal waters make up the deepest layers in the world ocean, even in temperate and tropical zones. 

Now, if you were to descend in a submarine to the depth where the mixed layer meets the abyssal waters and hold a thermometer out an open porthole, the submarine would fill with water, sink, and kill everyone aboard. (Words to live by: never open a porthole on a submerged submarine.) But if you attached a thermometer to the hull and connected it to a computer so that you could read the external water temperature as you descended, you would notice no change in temperature in the isothermal water column. But in a temperate ocean, you would reach a boundary between the winter-formed deep mixed layer and the polar-formed abyssal waters. At this boundary, you would notice a rapid change in temperature as you crossed from the bottom of the mixed layer to the top of the abyssal waters. This boundary—the region where temperature changes rapidly—is called a thermocline. The suffix -cline means “slope,” just like “incline” means up slope and “decline” means down slope. Where upper ocean waters—mixed ones—meet abyssal waters—ones formed in polar oceans—a permanent thermocline may be formed. Temporary thermoclines—seasonal ones—may be formed, too, as we’ll see in the next section.

The thermocline is a classic structure in descriptions of the physical properties of the ocean. It represents a boundary between different water parcels. The presence of a thermocline indicates that these water parcels are distinct and separate from each other. In graphs of the water column—XZ graphs with the z-axis representing depth and an x-axis representing temperature—the thermocline appears as a slope in the temperature profile. Thus, you may think of the thermocline as a temperature slope. 

13.7.2 Spring and the Seasonal Thermocline

As the days get longer and the sun gets higher in the sky in late winter or early spring (depending on the latitude), solar heating is more direct. The sea surface begins to warm. Air temperatures and water temperatures may be highly variable, subject to changing weather conditions. But at some point, the atmosphere in contact with the ocean becomes warmer than its surface. The net exchange of heat now is from the atmosphere to the ocean. When the sea surface warms, the seawater becomes less dense and more positively buoyant. But it doesn’t rise into the sky. Ocean water is about 800 times denser than even the most dense air, so the surface water remains at the surface.

The warming of the surface waters creates a temperature difference between the uppermost layer and the layer beneath it. The boundary between the now-warmer surface layer and the still-cold layer beneath it represents a different kind of thermocline, the seasonal thermocline. It forms as a result of surface warming and will disappear the following fall when the water cools again. It occurs seasonally—hence the name. 

Previous to the formation of the seasonal thermocline, the entire mixed layer is isothermal. Now the top of what used to be the mixed layer is warmer. The mixed layer has been split in two, so to speak. As a result, the mixed layer depth becomes shallower. The surface mixed layer is now confined to the region from the surface to the top of the seasonal thermocline. As the surface of the ocean continues to warm, multiple thermoclines may be present. It can get pretty complex, and we’re not going to dwell on all of the possibilities, but just be aware that the seasonal thermocline represents a temporary boundary between water parcels with different densities. The thermocline originates as the surface of the ocean warms. In a graph of the entire water column, from the surface to the seafloor, both the shallower seasonal thermocline and deeper permanent thermocline may be present. 

The layering of the ocean that occurs as a result of warming of surface waters and formation of a seasonal thermocline is called stratification. The word stratum means “layer,” and it’s used to describe clouds (e.g., stratus clouds), rocks (e.g., strata of sedimentary rock), or an archaeological dig (e.g., strata of ruins or refuse piles). Stratification of the water column signals a major event in the physics, chemistry, and biology of the upper ocean. It means, in effect, that the upper layers of the ocean are practically cut off from the lower layers. The upper and lower layers can’t mix, or mixing is severely restricted. Exchanges of heat, gases, salts, and biologically important nutrients between the surface mixed layer and lower layers can no longer occur or occur much more slowly. It won’t be until fall or even winter before these waters mix again. Keep this process in the back of your mind. When we get to chapters on productivity and food webs, the isolation of the upper ocean will help you appreciate how physical processes drive the chemistry and biology of the ocean.

13.7.3 Summer and the Multiple Thermoclines

The surface layer continues to warm well into summer as high sun angles permit solar radiation to penetrate deeper into the ocean. The upper ocean warms to its highest temperatures of the year. Double or even multiple thermoclines—more than a single thermocline visible in vertical profiles of the water column—may be observed (e.g., Fan et al. 2014). With layers of positively buoyant water parcels stacked on top of one another, the summer water column represents one of the most stable water columns of the year. The stability of the water column may lead to the formation of phytoplankton thin layers, bands of highly concentrated phytoplankton confined to a narrow depth interval. Having a vertical thickness of less than a few feet, they may extend horizontally for miles. Formation of thin layers requires a very stable water column, and in some places and under the right conditions, they may persist for weeks or longer (e.g., Durham and Stocker 2012).

Though plant life on land tends to be highly productive in summer, phytoplankton concentrations in temperate oceans tend to be low. That’s because stratification of surface waters limits mixing of biologically important nutrients. Starved of nutrients, phytoplankton grow more slowly. Species able to subsist under low nutrient concentrations replace those with high nutrient requirements. That’s just a snippet of the interaction of the ocean’s biology during the seasons of the sea.

13.7.4 Fall and the Disappearing Thermocline

In the temperate zone, all stable water columns must come to an end. The shortening of the days and the lower sun angles in fall reduce solar heating. The atmosphere cools quickly, much quicker than the ocean. At some point during fall, the air becomes cooler than the ocean, and the trend of heat exchange is reversed. The ocean begins to lose heat to the atmosphere. As the surface waters cool, they lose buoyancy and sink. Initially, they sink only a little. They remain plenty warm compared to the temperature minimum of winter, but it just takes a little cooling at the very surface of the ocean to make the water column unstable. As the cooling continues, the surface waters sink deeper. As the surface waters sink, they mix the upper ocean. Deeper waters begin to mix upward. Differences in heat, gases, salts, chemicals, biologically important nutrients, particles, plankton, and larvae disappear as the surface layer becomes colder and mixes deeper. The seasonal thermocline disappears too. The multiple layers of different density that were stacked up in the stable water column of summer become unstable and begin to blend and mix. The disappearance of the seasonal thermocline and the blending of the different layers characterize a process called destratification, the unmaking of the layers of the ocean. Of course, this process will continue throughout winter until the coldest days once again set the minimum temperature for the upper ocean.

13.8 The Importance of the Ocean’s Seasonal Cycle

The seasonal cycle of stratification and destratification drives some of the most important processes in the ocean. Seasonal production of cold water at the surface in polar oceans plays a key role in the formation of deep and abyssal waters, critical to the deep ocean circulation. The seasonal cycle of primary production in temperate and polar oceans nourishes marine food webs and supports fisheries. And the seasonal growth and decay of the polar ice caps has profound implications for a number of physical, chemical, and biological processes as well.

Formation of the seasonal thermocline in spring and summer and its disappearance in fall and winter represent one of the best examples of the connection between physical, chemical, and biological processes in the ocean. The seasonal dynamics of food webs—from tropical to polar ocean—emerge from the seasonal cycle of heating and cooling. In turn, these dynamics drive changes in the chemistry of the upper ocean, which have important implications for the carbon cycle and Earth’s climate. More will be revealed in the chapters ahead.

13.9 Chapter References

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Fan, Wei, Jinbao Song, and Shuang Li. 2014. “A Numerical Study on Seasonal Variations of the Thermocline in the South China Sea Based on the ROMS.” Acta Oceanologica Sinica 36: 56–64. https://doi.org/10.1007/s13131-014-0504-8

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