45 6.5 Climate Change
If one thing has been constant about Earth’s climate over geological time, it is its constant change. In the geological record, we can see this in the evidence of glaciations in the distant past, and we can also detect periods of extreme warmth by looking at the isotope composition of seafloor sediments. Not only has the climate changed frequently, the temperature fluctuations have been very significant. Today’s mean global temperature is about 15°C. However, during its coldest periods, the global mean was as cold as -50°C, while at various times during the Paleozoic and Mesozoic and during the Paleocene–Eocene thermal maximum, it was close to 30°C.
There are two parts to climate change, the first one is known as climate forcing, which is when conditions change to give the climate a little nudge in one direction or the other. The second part of climate change, and the one that typically does most of the work, is what we call a feedback. When a climate forcing changes the climate a little, a whole series of environmental changes take place, many of which either exaggerate the initial change (positive feedback), or suppress the change (negative feedback).
An example of a climate forcing mechanism is the increase in the amount of carbon dioxide (CO2) in the atmosphere that results from our use of fossil fuels. CO2 traps heat in the atmosphere and leads to climate warming. Warming changes vegetation patterns; contributes to the melting of snow, ice, and permafrost; causes sea level to rise; reduces the solubility of CO2 in sea water; and has a number of other minor effects. Most of these changes contribute to more warming. Melting of permafrost, for example, is a strong positive feedback because frozen soil contains trapped organic matter that is converted to CO2 and methane (CH4) when the soil thaws. Both these gases accumulate in the atmosphere and add to the warming effect. On the other hand, if warming causes more vegetation growth, that vegetation should absorb CO2, thus reducing the warming effect, which would be a negative feedback. Under our current conditions — a planet that still has lots of glacial ice and permafrost — most of the feedbacks that result from a warming climate are positive feedbacks and so the climate changes that we cause get naturally amplified by natural processes.
Natural Climate Forcing
Natural climate forcing has been going on throughout geological time. A wide range of processes has been operating at widely different time scales, from a few years to billions of years. The longest-term natural forcing variation is related to the evolution of the Sun. Like most other stars of a similar mass, our Sun is evolving. For the past 4.6 billion years, its rate of nuclear fusion has been increasing, and it is now emitting about 40% more energy (as light) than it did at the beginning of geological time. A difference of 40% is big, so it’s a little surprising that the temperature on Earth has remained at a reasonable and habitable temperature for all of this time. The mechanism for that relative climate stability has been the evolution of our atmosphere from one that was dominated by CO2, and also had significant levels of CH4 — both greenhouse gasses — to one with only a few hundred parts per million of CO2 and just under 1 part per million of CH4. Those changes to our atmosphere have been no accident; over geological time, life and its metabolic processes have evolved (such as the evolution of photosynthetic bacteria that consume CO2) and changed the atmosphere to conditions that remained cool enough to be habitable.
The position of the Earth relative to the Sun is another important component of natural climate forcing. Earth’s orbit around the Sun is nearly circular, but like all physical systems, it has natural oscillations. First, the shape of the orbit changes on a regular time scale (close to 100,000 years) from being close to circular to being very slightly elliptical. But the circularity of the orbit is not what matters; it is the fact that as the orbit becomes more elliptical, the position of the Sun within that ellipse becomes less central or more eccentric (Figure 6.5.1a). Eccentricity is important because when it is high, the Earth-Sun distance varies more from season to season than it does when eccentricity is low.
Second, Earth rotates around an axis through the North and South Poles, and that axis is at an angle to the plane of Earth’s orbit around the Sun (Figure 6.5.1b). The angle of tilt (also known as obliquity) varies on a time scale of 41,000 years. When the angle is at its maximum (24.5°), Earth’s seasonal differences are accentuated. When the angle is at its minimum (22.1°), seasonal differences are minimized. The current hypothesis is that glaciation is favored at low seasonal differences as summers would be cooler and snow would be less likely to melt and more likely to accumulate from year to year. Third, the direction in which Earth’s rotational axis points also varies, on a time scale of about 20,000 years (Figure 6.5.1c). This variation, known as precession, means that although the North Pole is presently pointing to the star Polaris (the pole star), in 10,000 years it will point to the star Vega. The importance of eccentricity, tilt, and precession to Earth’s climate cycles (now known as Milankovitch Cycles) was first pointed out by Yugoslavian engineer and mathematician Milutin Milankovitch in the early 1900s. Milankovitch recognized that although the variations in the orbital cycles did not affect the total amount of insolation (light energy from the Sun) that Earth received, it did affect where on Earth that energy was strongest.
Volcanic eruptions don’t just involve lava flows and exploding rock fragments; various particulates and gases are also released, the important ones being sulphur dioxide and CO2. Sulphur dioxide is an aerosol that reflects incoming solar radiation and has a net cooling effect that is short lived (a few years in most cases, as the particulates settle out of the atmosphere within a couple of years), and doesn’t typically contribute to longer-term climate change. Volcanic CO2 emissions can contribute to climate warming but only if a greater-than-average level of volcanism is sustained over a long time (at least tens of thousands of years). It is widely believed that the catastrophic end-Permian extinction (at 250 Ma) resulted from warming initiated by the eruption of the massive Siberian Traps over a period of at least a million years.
Ocean currents are important to climate, and currents also have a tendency to oscillate. Glacial ice cores show clear evidence of changes in the Gulf Stream that affected global climate on a time scale of about 1,500 years during the last glaciation. The east-west changes in sea-surface temperature and surface pressure in the equatorial Pacific Ocean, known as the El Niño Southern Oscillation or ENSO varies on a much shorter time scale of between two and seven years. These variations tend to garner the attention of the public because they have significant climate implications in many parts of the world. The strongest El Niños in recent decades were in 1983, 1998, and 2015 and those were very warm years from a global perspective. During a strong El Niño, the equatorial Pacific sea-surface temperatures are warmer than normal and heat the atmosphere above the ocean, which leads to warmer-than-average global temperatures.
Climate Feedbacks
As already stated, climate feedbacks are critically important in amplifying weak climate forcings into full-blown climate changes. Since Earth still has a very large volume of ice, mostly in the continental ice sheets of Antarctica and Greenland, but also in alpine glaciers and permafrost, melting is one of the key feedback mechanisms. Melting of ice and snow leads to several different types of feedbacks, an important one being a change in albedo, or the reflectivity of a surface. Earth’s various surfaces have widely differing albedos, expressed as the percentage of light that reflects off a given material. This is important because most solar energy that hits a very reflective surface is not absorbed and therefore does little to warm Earth. Water in the oceans or on a lake is one of the darkest surfaces, reflecting less than 10% of the incident light, while clouds and snow or ice are among the brightest surfaces, reflecting 70% to 90% of the incident light. When sea ice melts, as it has done in the Arctic Ocean at a disturbing rate over the past decade, the albedo of the area affected changes dramatically, from around 80% down to less than 10%. Much more solar energy is absorbed by the water than by the pre-existing ice, and the temperature increase is amplified. The same applies to ice and snow on land, but the difference in albedo is not as great. When ice and snow on land melt, sea level rises. (Sea level is also rising because the oceans are warming and that increases their volume). A higher sea level means a larger proportion of the planet is covered with water, and since water has a lower albedo than land, more heat is absorbed and the temperature goes up a little more. Since the last glaciation, sea-level rise has been about 125 m; a huge area that used to be land is now flooded by heat-absorbent seawater. During the current period of anthropogenic climate change, sea level has risen only about 20 cm, and although that doesn’t make a big change to albedo, sea-level rise is accelerating.
Most of northern Canada, Alaska, Russia, and Scandinavia has a layer of permafrost that ranges from a few centimeters to hundreds of meters in thickness. Permafrost is a mixture of soil and ice and it also contains a significant amount of trapped organic carbon that is released as CO2 and CH4 when the permafrost breaks down. Because the amount of carbon stored in permafrost is in the same order of magnitude as the amount released by burning fossil fuels, this is a feedback mechanism that has the potential to equal or surpass the forcing that has unleashed it. In some polar regions, including northern Canada, permafrost includes methane hydrate, a highly concentrated form of CH4 trapped in solid form. Breakdown of permafrost releases this CH4. Even larger reserves of methane hydrate exist on the seafloor, and while it would take significant warming of ocean water down to a depth of hundreds of meters, this too is likely to happen in the future if we don’t limit our impact on the climate. There is strong isotopic evidence that the Paleocene–Eocene thermal maximum was caused, at least in part, by a massive release of sea-floor methane hydrate.
There is about 45 times as much carbon in the ocean (as dissolved bicarbonate ions, HCO3–) as there is in the atmosphere (as CO2), and there is a steady exchange of carbon between the two reservoirs (see section 5.5). But the solubility of CO2 in water decreases as the temperature goes up. In other words, the warmer it gets, the more oceanic bicarbonate that gets transferred to the atmosphere as CO2. That makes CO2 solubility another positive feedback mechanism. Vegetation growth responds positively to both increased temperatures and elevated CO2 levels, and so in general, it represents a negative feedback to climate change because the more the vegetation grows, the more CO2 is taken from the atmosphere. But it’s not quite that simple, because when trees grow bigger and more vigorously, forests become darker (they have lower albedo) so they absorb more heat. Furthermore, climate warming isn’t necessarily good for vegetation growth; some areas have become too hot, too dry, or even too wet to support the plant community that was growing there, and it might take centuries for something to replace it successfully. All of these positive (and negative) feedbacks work both ways. For example, during climate cooling, growth of glaciers leads to higher albedos, and formation of permafrost results in storage of carbon that would otherwise have returned quickly to the atmosphere.
Anthropogenic Climate Change
When we talk about anthropogenic climate change, we are generally thinking of the industrial era, which really got going when we started using fossil fuels (coal to begin with, and later oil and natural gas) to drive machinery and trains, and to generate electricity. That was around the middle of the 18th century. The issue with fossil fuels is that they involve burning carbon that was naturally stored in the crust over hundreds of millions of years as part of Earth’s process of counteracting the warming Sun.
A rapidly rising population, the escalating level of industrialization and mechanization of our lives, and an increasing dependence on fossil fuels have driven the anthropogenic climate change of the past century. The trend of mean global temperatures since 1850 is shown in Figure 6.5.2. For approximately the past 55 years, the temperature has increased at a relatively steady and disturbingly rapid rate, especially compared to past changes. The average temperature now is approximately 1.1°C higher than before industrialization, and two-thirds of this warming has occurred since 1975.
The Intergovernmental Panel on Climate Change (IPCC), established by the United Nations in 1988, is responsible for reviewing the scientific literature on climate change and issuing periodic reports on several topics, including the scientific basis for understanding climate change, our vulnerability to observed and predicted climate changes, and what we can do to limit climate change and minimize its impacts. Figure 6.5.3, from the sixth report of the IPCC, issued in preliminary form in 2021, shows the relative contributions of various greenhouse gases and other factors to current climate forcing, based on the changes from levels that existed in 1750.
The biggest anthropogenic contributor to warming is the emission of CO2, which accounts for 50% of positive forcing. CH4 and its atmospheric derivatives (CO2, H2O, and O3) account for 29%, and the halocarbon gases (mostly leaked from air-conditioning appliances) and nitrous oxide (N2O) (from burning fossils fuels) account for 5% each. Carbon monoxide (CO) (also produced by burning fossil fuels) accounts for 7%, and the volatile organic compounds other than methane (NMVOC) account for 3%. CO2 emissions come mostly from coal- and gas-fired power stations, motorized vehicles (cars, trucks, and aircraft), and industrial operations (e.g., smelting), and indirectly from forestry. CH4 emissions come from production of fossil fuels (escape from coal mining and from gas and oil production), livestock farming (mostly beef), landfills, and wetland rice farming. N2O and CO come mostly from the combustion of fossil fuels. In summary, close to 70% of our current greenhouse gas emissions come from fossil fuel production and use, while most of the rest comes from agriculture and landfills. Figure 6.5.4 shows the IPCC’s projections for temperature increases over the next 100 years as a result of these increasing greenhouse gases.
Impacts of Climate Change
We’ve all experienced the effects of climate change over the past decade. However, it’s not straightforward for climatologists to make the connection between a warming climate and specific weather events, and most are justifiably reluctant to ascribe any specific event to climate change. In this respect, the best measures of climate change are those that we can detect over several decades, such as the temperature changes shown in Figure 6.5.2, or the sea level rise shown in Figure 6.5.5. As already stated, sea level has risen approximately 20 cm since 1750, and that rise is attributed to both warming (and therefore expanding) seawater and melting glaciers and other land-based snow and ice (melting of sea ice does not contribute directly to sea level rise as it is already floating in the ocean).
Projections for sea level rise to the end of this century vary widely. This is in large part because we do not know which of the above climate change scenarios (Figure 6.5.4) we will most closely follow, but many are in the range from 0.5 m to 2.0 m. One of the problems in predicting sea level rise is that we do not have a strong understanding of how large ice sheets, such as Greenland and Antarctica, will respond to future warming. Another issue is that the oceans don’t respond immediately to warming. For example, with the current amount of warming, we are already committed to a future sea level rise of between 1.3 m and 1.9 m, even if we could stop climate change today. This is because it takes decades to centuries for the existing warming of the atmosphere to be transmitted to depth within the oceans and to exert its full impact on large glaciers. Most of that committed rise would take place over the next century, but some would be delayed longer. And for every decade that the current rates of climate change continue, that number increases by another 0.3 m. In other words, if we don’t make changes quickly, by the end of this century we’ll be locked into 3 m of future sea level rise. In a 2008 report, the Organization for Economic Co-operation and Development (OECD) estimated that by 2070 approximately 150 million people living in coastal areas could be at risk of flooding due to the combined effects of sea level rise, increased storm intensity, and land subsidence. The assets at risk (buildings, roads, bridges, ports, etc.) are in the order of $35 trillion ($35,000,000,000,000). Countries with the greatest exposure of population to flooding are China, India, Bangladesh, Vietnam, U.S.A., Japan, and Thailand. Some of the major cities at risk include Shanghai, Guangzhou, Mumbai, Kolkata, Dhaka, Ho Chi Minh City, Tokyo, Miami, and New York.
One of the other risks for coastal populations, besides sea level rise, is that climate warming is also associated with an increase in the intensity of tropical storms (e.g., hurricanes or typhoons; see section 6.4), which almost always bring serious flooding from intense rain and storm surges. Some recent examples are New Orleans in 2005 with Hurricane Katrina, and New Jersey and New York in 2012 with Hurricane Sandy. Tropical storms get their energy from the evaporation of warm seawater in tropical regions. In the Atlantic Ocean, this takes place between 8° and 20° N in the summer. Figure 6.5.6 shows the variations in the sea-surface temperature (SST) of the tropical Atlantic Ocean (in blue) versus the amount of power represented by Atlantic hurricanes between 1950 and 2008 (in red). Not only has the overall intensity of Atlantic hurricanes increased with the warming since 1975, but the correlation between hurricanes and sea-surface temperatures is very strong over that time period.
The geographical ranges of diseases and pests, especially those caused or transmitted by insects, have been shown to extend toward temperate regions because of climate change. West Nile virus and Lyme disease are two examples that already directly affect North Americans, while dengue fever could be an issue in the future (dengue became a “nationally notifiable condition” in the United States in 2010). For several weeks in July and August of 2010, a massive heat wave affected western Russia, especially the area southeast of Moscow, and scientists have stated that climate change was a contributing factor. Temperatures soared to over 40°C, as much as 12°C above normal over a wide area, and wildfires raged in many parts of the country. Over 55,000 deaths are attributed to the heat and to respiratory problems associated with the fires. A summary of the impacts of climate change on natural disasters is given in Figure 6.5.7. The major types of disasters related to climate are floods and storms, but the health implications of extreme temperatures are also becoming a great concern. In the decade 1971 to 1980, extreme temperatures were the fifth most common natural disasters; by 2001 to 2010, they were the third most common.
Oxygen and carbon dioxide are involved in the same biological processes in the ocean, but in opposite ways; photosynthesis consumes CO2 and produces O2, while respiration and decomposition consume O2 and produce CO2. Therefore it should not be surprising that oceanic CO2 profiles are essentially the opposite of dissolved oxygen profiles (Figure 5.5.1). At the surface, photosynthesis consumes CO2 so CO2 levels remain relatively low. In addition, organisms that utilize carbonate in their shells are common near the surface, further reducing the amount of dissolved CO2.
In deeper water, CO2 concentration increases as respiration exceeds photosynthesis, and decomposition of organic matter adds additional CO2 to the water. As with oxygen, there is often more CO2 at depth because cold bottom water holds more dissolved gases, and high pressures increase solubility. Deep water in the Pacific contains more CO2 than the Atlantic as the Pacific water is older and has accumulated more CO2 from the respiration of benthic organisms.
But the behavior of carbon dioxide in the ocean is more complex than the figure above would suggest. When CO2 gas dissolves in the ocean, it interacts with the water to produce a number of different compounds according to the reaction below:
CO2 + H2O ↔ H2CO3 ↔ H+ + HCO3- ↔ 2H+ + CO32-
CO2 reacts with water to produce carbonic acid (H2CO3), which then dissociates into bicarbonate (HCO3-) and hydrogen ions (H+). The bicarbonate ions can further dissociate into carbonate (CO32-) and additional hydrogen ions (Figure 5.5.2).
Most of the CO2 dissolving or produced in the ocean is quickly converted to bicarbonate. Bicarbonate accounts for about 92% of the CO2 dissolved in the ocean, and carbonate represents around 7%, so only about 1% remains as CO2, and little gets absorbed back into the air. The rapid conversion of CO2 into other forms prevents it from reaching equilibrium with the atmosphere, and in this way, water can hold 50-60 times as much CO2 and its derivatives as the air.
CO2 and pH
The equation above also illustrates carbon dioxide's role as a buffer, regulating the pH of the ocean. Recall that pH reflects the acidity or basicity of a solution. The pH scale runs from 0-14, with 0 indicating a very strong acid, and 14 representing highly basic conditions. A solution with a pH of 7 is considered neutral, as is the case for pure water. The pH value is calculated as the negative logarithm of the hydrogen ion concentration according to the equation:
pH = -log10[H+]
Therefore, a high concentration of H+ ions leads to a low pH and acidic condition, while a low H+ concentration indicates a high pH and basic conditions. It should also be noted that pH is described on a logarithmic scale, so every one point change on the pH scale actually represents an order of magnitude (10 x) change in solution strength. So a pH of 6 is 10 times more acidic than a pH of 7, and a pH of 5 is 100 times (10 x 10) more acidic than a pH of 7.
Carbon dioxide and the other carbon compounds listed above play an important role in buffering the pH of the ocean. Currently, the average pH for the global ocean is about 8.1, meaning seawater is slightly basic. Because most of the inorganic carbon dissolved in the ocean exists in the form of bicarbonate, bicarbonate can respond to disturbances in pH by releasing or incorporating hydrogen ions into the various carbon compounds. If pH rises (low [H+]), bicarbonate may dissociate into carbonate, and release more H+ ions, thus lowering pH. Conversely, if pH gets too low (high [H+]), bicarbonate and carbonate may incorporate some of those H+ ions and produce bicarbonate, carbonic acid, or CO2 to remove H+ ions and raise the pH. By shuttling H+ ions back and forth between the various compounds in this equation, the pH of the ocean is regulated and conditions remain favorable for life.
CO2 and Ocean Acidification
In recent years there has been rising concern about the phenomenon of ocean acidification. As described in the processes above, the addition of CO2 to seawater lowers the pH of the water. As anthropogenic sources of atmospheric CO2 have increased since the Industrial Revolution, the oceans have been absorbing an increasing amount of CO2, and researchers have documented a decline in ocean pH from about 8.2 to 8.1 in the last century. This may not appear to be much of a change, but remember that since pH is on a logarithmic scale, this decline represents a 30% increase in acidity. It should be noted that even at a pH of 8.1 the ocean is not actually acidic; the term "acidification" refers to the fact that the pH is becoming lower, i.e. the water is moving towards more acidic conditions.
Figure 5.5.3 presents data from observation stations in and around the Hawaiian Islands. As atmospheric levels of CO2 have increased, the CO2 content of the ocean water has also increased, leading to a reduction in seawater pH. Some models suggest that at the current rate of CO2 addition to the atmosphere, by 2100 ocean pH may be further reduced to around 7.8, which would represent more than a 120% increase in ocean acidity since the Industrial Revolution.
Why is this important? Declining pH can impact many biological systems. Of particular concern are organisms that secrete calcium carbonate shells or skeletons, such as corals, shellfish, and may planktonic organisms. At lower pH levels, calcium carbonate dissolves, eroding the shells and skeletons of these organisms (Figure 5.5.4).
Not only does a declining pH lead to increased rates of dissolution of calcium carbonate, it also diminishes the amount of free carbonate ions in the water. The relative proportions of the different carbon compounds in seawater is dependent on pH (Figure 5.5.6). As pH declines, the amount of carbonate declines, so there is less available for organisms to incorporate into their shells and skeletons. So ocean acidification both dissolves existing shells and makes it harder for shell formation to occur.
Additional links for more information:
- NOAA Ocean Acidification Program website http://oceanacidification.noaa.gov/
The primary surface current along the east coast of the United States is the Gulf Stream, which was first mapped by Benjamin Franklin in the 18th century (Figure 7.2.1). As a strong, fast current, it reduced the sailing time for ships traveling from the United States back to Europe, so sailors would use thermometers to locate its warm water and stay within the current.
The Gulf Stream is formed from the convergence of the North Atlantic Equatorial Current bringing tropical water from the east, and the Florida Current that brings warm water from the Gulf of Mexico. The Gulf Stream takes this warm water and transports it northwards along the U.S. east coast (Figure 7.2.2). As a western boundary current, the Gulf Stream experiences western intensification (section 7.4), making the current narrow (50-100 km wide), deep (to depths of 1.5 km) and fast. With an average speed of 6.4 km/hr, and a maximum speed of about 9 km/hr, it is the fastest current in the world ocean. It also transports huge amounts of water, more than 100 times greater than the combined flow of all of the rivers on Earth.
As the Gulf Stream approaches Canada, the current becomes wider and slower as the flow dissipates and it encounters the cold Labrador Current moving in from the north. At this point, the current begins to meander, or change from a fast, straight flow to a slower, looping current (Figure 7.2.2). Often these meanders loop so much that they pinch off and form large rotating water masses called rings or eddies, that separate from the Gulf Stream. If an eddy pinches off from the north side of the Gulf Stream, it entraps a mass of warm water and moves it north into the surrounding cold water of the North Atlantic. These warm core rings are shallow, bowl-shaped water masses about 1 km deep, and about 100 km across, that rotate clockwise as they carry warm water in to the North Atlantic (Figure 7.2.3). If the meanders pinch off at the southern boundary of the Gulf Stream, they form cold core rings that rotate counterclockwise and move to the south. Cold core rings are cone-shaped water masses extending down to over 3.5 km deep, and may be over 500 km wide at the surface.
After the Gulf Stream meets the cold Labrador Current, it joins the North Atlantic Current, which transports the warm water towards Europe, where it moderates the European climate. It is estimated that Northern Europe is up to 9o C warmer than expected because of the Gulf Stream, and the warm water helps to keep many northern European ports ice-free in the winter.
In the east, the Gulf Stream merges into the Sargasso Sea, which is the area of the ocean within the rotation center of the North Atlantic gyre. The Sargasso Sea gets its name from the large floating mats of the marine algae Sargassum that are abundant on the surface (Figure 7.2.4). These Sargassum mats may play an important role in the early life stages of sea turtles, who may live and feed within the algae for many years before reaching adulthood.
The primary surface current along the east coast of the United States is the Gulf Stream, which was first mapped by Benjamin Franklin in the 18th century (Figure 7.2.1). As a strong, fast current, it reduced the sailing time for ships traveling from the United States back to Europe, so sailors would use thermometers to locate its warm water and stay within the current.
The Gulf Stream is formed from the convergence of the North Atlantic Equatorial Current bringing tropical water from the east, and the Florida Current that brings warm water from the Gulf of Mexico. The Gulf Stream takes this warm water and transports it northwards along the U.S. east coast (Figure 7.2.2). As a western boundary current, the Gulf Stream experiences western intensification (section 7.4), making the current narrow (50-100 km wide), deep (to depths of 1.5 km) and fast. With an average speed of 6.4 km/hr, and a maximum speed of about 9 km/hr, it is the fastest current in the world ocean. It also transports huge amounts of water, more than 100 times greater than the combined flow of all of the rivers on Earth.
As the Gulf Stream approaches Canada, the current becomes wider and slower as the flow dissipates and it encounters the cold Labrador Current moving in from the north. At this point, the current begins to meander, or change from a fast, straight flow to a slower, looping current (Figure 7.2.2). Often these meanders loop so much that they pinch off and form large rotating water masses called rings or eddies, that separate from the Gulf Stream. If an eddy pinches off from the north side of the Gulf Stream, it entraps a mass of warm water and moves it north into the surrounding cold water of the North Atlantic. These warm core rings are shallow, bowl-shaped water masses about 1 km deep, and about 100 km across, that rotate clockwise as they carry warm water in to the North Atlantic (Figure 7.2.3). If the meanders pinch off at the southern boundary of the Gulf Stream, they form cold core rings that rotate counterclockwise and move to the south. Cold core rings are cone-shaped water masses extending down to over 3.5 km deep, and may be over 500 km wide at the surface.
After the Gulf Stream meets the cold Labrador Current, it joins the North Atlantic Current, which transports the warm water towards Europe, where it moderates the European climate. It is estimated that Northern Europe is up to 9o C warmer than expected because of the Gulf Stream, and the warm water helps to keep many northern European ports ice-free in the winter.
In the east, the Gulf Stream merges into the Sargasso Sea, which is the area of the ocean within the rotation center of the North Atlantic gyre. The Sargasso Sea gets its name from the large floating mats of the marine algae Sargassum that are abundant on the surface (Figure 7.2.4). These Sargassum mats may play an important role in the early life stages of sea turtles, who may live and feed within the algae for many years before reaching adulthood.
The most obvious feature of the oceans is that they contain water. Water is so ubiquitous that it may not seem like a very interesting substance, but it has many unique properties that impact global oceanographic and climatological processes. Many of these processes are due to hydrogen bonds forming between water molecules.
The water molecule consists of two hydrogen atoms and one oxygen atom. The electrons responsible for the bonds between the atoms are not distributed equally throughout the molecule, so that the hydrogen ends of water molecules have a slight positive charge, and the oxygen end has a slight negative charge, making water a polar molecule. The negative oxygen side of the molecule forms an attraction to the positive hydrogen end of a neighboring molecule. This rather weak force of attraction is called a hydrogen bond (Figure 5.1.1). If not for hydrogen bonds, water would vaporize at -68o C, meaning liquid water (and thus life) could not exist on Earth. These hydrogen bonds are responsible for some of water’s unique properties:
1. Water is the only substance to naturally exist in a solid, liquid, and gaseous form under the normal range of temperatures and pressures found on Earth. This is due to water’s relatively high freezing and vaporizing points (see below).
2. Water has a high heat capacity, which is the amount of heat that must be added to raise its temperature. Specific heat is the heat required to raise the temperature of 1 g of a substance by 1o C. Water has the highest specific heat of any liquid except ammonia (Table 5.1.1).
Table 5.1.1 Specific heat values for a number of common substances
| Specific Heat (calories/g/Co) | |
|---|---|
| Ammonia | 1.13 |
| Water | 1.00 |
| Acetone | 0.51 |
| Grain Alcohol | 0.23 |
| Aluminum | 0.22 |
| Copper | 0.09 |
| Silver | 0.06 |
Water is therefore one of the most difficult liquids to heat or cool; it can absorb large amounts of heat without increasing its temperature. Remember that temperature reflects the average kinetic energy of the molecules within a substance; the more vigorous the motion, the higher the temperature. In water, the molecules are held together by hydrogen bonds, and these bonds must be overcome to allow the molecules to move freely. When heat is added to water the energy must first go to breaking the hydrogen bonds before the temperature can begin to rise. Therefore, much of the added heat is absorbed by breaking H bonds, not by increasing the temperature, giving water a high heat capacity.
Hydrogen bonds also give water a high latent heat; the heat required to undergo a phase change from solid to liquid, or liquid to gas. The latent heat of fusion is the heat required to go from solid to liquid; 80 cal/g in the case of ice melting to water. Ice is a solid because hydrogen bonds hold the water molecules into a solid crystal lattice (see below). As ice is heated, the temperature rises up to 0o C. At that point, any additional heat goes to melting the ice by breaking the hydrogen bonds, not to increasing the temperature. So as long as ice is present, the water temperature will not increase. This is why your drink will remain cold as long as it contains ice; any heat absorbed goes to melting the ice, not to warming the drink.
When all of the ice is melted, additional heat will increase the temperature of the water 1o C for each calorie of heat added, until it reaches 100o C. At that point, any additional heat goes to overcoming the hydrogen bonds and turning the liquid water into water vapor, rather than increasing the water temperature. The heat required to evaporate liquid water into water vapor is the latent heat of vaporization which has a value of 540 cal/g (Figure 5.1.2).
The high heat capacity of water helps regulate global climate, as the oceans slowly absorb and release heat, preventing rapid swings in temperature (see section 8.1). It also means that aquatic organisms aren't as subjected to the same rapid temperature changes as terrestrial organisms. A deep ocean organism may not experience more than a 0.5o C change in temperature over its entire life, while a terrestrial species may encounter changes of more than 20o C in a single day!
3. Water dissolves more substances than any other liquid; it is a "universal solvent", which is why so many substances are dissolved in the ocean. Water is especially good at dissolving ionic salts; molecules made from oppositely charged ions such as NaCl (Na+ and Cl-). In water, the charged ions attract the polar water molecules. The ions get surrounded by a layer of water molecules, weakening the bond between the ions by up to 80 times. With the bonds weakened between ions, the substance dissolves (Figure 5.1.3).
4. The solid phase is less dense than the liquid phase. In other words, ice floats. Most substances are denser in the solid form than in the liquid form, as their molecules are more closely packed together as a solid. Water is an exception: the density of fresh water is 1.0 g/cm3, while the density of ice is 0.92 g/cm3, and once again, this is due to the action of hydrogen bonds.
As water temperature cools the molecules slow down, eventually slowing enough that hydrogen bonds can form and hold the water molecules in a crystal lattice. The molecules in the lattice are spaced farther apart than the molecules in liquid water, which makes ice less dense than liquid water (Figure 5.1.4). This is familiar to anyone who has ever left a full water bottle in the freezer, only to have it burst as the water freezes and expands.
But the relationship between temperature and water density is not a simple linear one. As water cools, its density increases as expected, as the water molecules slow down and get closer together. However, fresh water reaches its maximum density at a temperature of 4o C, and as it cools beyond that point its density declines as the hydrogen bonds begin to form and the intermolecular spacing increases (Figure 5.1.5 inset). The density continues to decline until the temperature reaches 0o C and ice crystals form, reducing the density dramatically (Figure 5.1.5).
There are a number of important implications to ice being less dense than water. Ice floating on the surface of the ocean helps regulate ocean temperatures, and therefore global climate, by influencing the amount of sunlight that is reflected rather than absorbed (see section 5.6). On a smaller scale, surface ice can prevent lakes and ponds from freezing solid during the winter. As fresh surface water cools, the water gets denser, and sinks to the bottom. The new surface water then cools and sinks, and the process is repeated in what is referred to as overturning, with denser water sinking and less dense water moving to the surface only to be cooled and sink itself. In this way, the entire body of water is cooled somewhat evenly. This process continues until the surface water cools below 4o C. Below 4o C, the water becomes less dense as it cools, so it no longer sinks. Instead, it remains as the surface, getting colder and less dense, until it freezes at 0o C. Once fresh water freezes, the ice floats and insulates the rest of the water beneath it, reducing further cooling. The densest bottom water is still at 4o C, so it does not freeze, allowing the bottom of a lake or pond to remain unfrozen (which is good news for the animals living there) no matter how cold it gets outside.
The dissolved salts in seawater inhibit the formation of the crystal lattice, and therefore make it harder for ice to form. So seawater has a freezing point of about -2o C (depending on salinity), and freezes before a temperature of maximum density is reached. Thus seawater will continue to sink as it gets colder, until it finally freezes.
5. Water has a very high surface tension, the highest of any liquid except mercury (Table 5.1.2). Water molecules are attracted to each other by hydrogen bonds. For molecules not at the water surface, they are surrounded by other water molecules in all directions, so the attractive forces are evenly distributed in all directions. But for molecules at the surface there are few adjacent molecules above them, only below, so all of the attractive forces are directed inwards, away from the surface (Figure 5.1.6). This inwards force is what causes water droplets to take on a spherical shape, and water to bead up on a surface, as the spherical shape provides the minimum possible surface area. These attractive forces also cause the surface of the water to act like an elastic "skin" which allows things like insects to sit on the water's surface without sinking.
Table 5.1.2 Surface tensions of various liquids
| Liquid | Surface Tension (millinewton/meter) | Temperature oC |
|---|---|---|
| Mercury | 487.00 | 15 |
| Water | 71.97 | 25 |
| Glycerol | 63.00 | 20 |
| Acetone | 23.70 | 20 |
| Ethanol | 22.27 | 20 |
By Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
Most of the waves discussed in the previous section referred to deep water waves in the open ocean. But what happens when these waves move towards shore and encounter shallow water? Remember that in deep water, a wave’s speed depends on its wavelength, but in shallow water wave speed depends on the depth (section 3.1). When waves approach the shore they will "touch bottom" at a depth equal to half of their wavelength; in other words, when the water depth equals the depth of the wave base (Figure 3.3.1). At this point their behavior will begin to be influenced by the bottom.
When the wave touches the bottom, friction causes the wave to slow down. As one wave slows down, the one behind it catches up to it, thus decreasing the wavelength. However, the wave still contains the same amount of energy, so while the wavelength decreases, the wave height increases. Eventually the wave height exceeds 1/7 of the wavelength, and the wave becomes unstable and forms a breaker. Often breakers will start to curl forwards as they break. This is because the bottom of the wave begins to slow down before the top of the wave, as it is the first part to encounter the seafloor. So the crest of the wave gets “ahead” of the rest of the wave, but has no water underneath it to support it (Figure 3.3.1).
There are three main types of breakers: spilling, plunging, and surging. These are related to the steepness of the bottom, and how quickly the wave will slow down and its energy will get dissipated.
- Spilling breakers form on gently sloping or flatter beaches, where the energy of the wave is dissipated gradually. The wave slowly increases in height, then slowly collapses on itself (Figure 3.3.2). For surfers, these waves provide a longer ride, but they are less exciting.
- Plunging breakers form on more steeply-sloped shores, where there is a sudden slowing of the wave and the wave gets higher very quickly. The crest outruns the rest of the wave, curls forwards and breaks with a sudden loss of energy (Figure 3.3.3). These are the “pipeline” waves that surfers seek out.
- Surging breakers form on the steepest shorelines. The wave energy is compressed very suddenly right at the shoreline, and the wave breaks right onto the beach (Figure 3.3.4). These waves give too short (and potentially painful) a ride for surfers to enjoy.
Wave Refraction
Swell can be generated anywhere in the ocean and therefore can arrive at a beach from almost any direction. But if you have ever stood at the shore you have probably noticed that the waves usually approach the shore somewhat parallel to the coast. This is due to wave refraction. If a wave front approaches shore at an angle, the end of the wave front closest to shore will touch bottom before the rest of the wave. This will cause that shallower part of the wave to slow down first, while the rest of the wave that is still in deeper water will continue on at its regular speed. As more and more of the wave front encounters shallower water and slows down, the wave font refracts and the waves tend to align themselves nearly parallel to the shoreline (they are refracted towards the region of slower speed). As we will see in section 5.2, the fact that the waves do not arrive perfectly parallel to the beach causes longshore currents and longshore transport that run parallel to the shore.
Refraction can also explain why waves tend to be larger off of points and headlands, and smaller in bays. A wave front approaching shore will touch the bottom off of the point before it touches bottom in a bay. Once again, the shallower part of the wave front will slow down, and cause the rest of the wave front to refract towards the slower region (the point). Now all of the initial wave energy is concentrated in a relatively small area off of the point, creating large, high energy waves (Figure 3.3.6). In the bay, the refraction has caused the wave fronts to refract away from each other, dispersing the wave energy, and leading to calmer water and smaller waves. This makes the large waves of a “point break” ideal for surfing, while water is calmer in a bay, which is where people would launch a boat. This difference in wave energy also explains why there is net erosion on points, while sand and sediments get deposited in bays (see section 5.3).
By Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
Estuaries are partially enclosed bodies of water where the salt water is diluted by fresh water input from land, creating brackish water with a salinity somewhere between fresh water and normal seawater. Estuaries include many bays, inlets, and sounds, and are often subject to large temperature and salinity variations due to their enclosed nature and smaller size compared to the open ocean.
Estuaries can be classified geologically into four basic categories based on their method of origin. In all cases they are a result of rising sea level over the last 18,000 years, beginning with the end of the last ice age; a period that has seen a rise of about 130 m. The rise in sea level has flooded coastal areas that were previously above water, and prevented the estuaries from being filled in by all of the sediments that have been emptied into them.
The first type is a coastal plain estuary, or drowned river valley. These estuaries are formed as sea level rises and floods an existing river valley, mixing salt and fresh water to create the brackish conditions where the river meets the sea. These types of estuaries are common along the east coast of the United States, including major bodies such as the Chesapeake Bay, Delaware Bay, and Narragansett Bay (Figure 4.6.1). Coastal plain estuaries are usually shallow, and since there is a lot of sediment input from the rivers, there are often a number of depositional features associated with them such as spits and barrier islands.
The presence of sand bars, spits, and barrier islands can lead to bar-built estuaries, where a barrier is created between the mainland and the ocean. The water that remains inside the sand bar is cut off from complete mixing with the ocean, and receives freshwater input from the mainland, creating estuarine conditions (Figure 4.6.2).
Fjords are estuaries formed in deep, U-shaped basins that were carved out by advancing glaciers. When the glaciers melted and retreated, sea level rose and filled these troughs, creating deep, steep-walled fjords (Figure 4.6.3). Fjords are common in Norway, Alaska, Canada, and New Zealand, where there are mountainous coastlines once covered by glaciers.
Tectonic estuaries are the result of tectonic movements, where faulting causes some sections of the crust to subside, and those lower elevation sections then get flooded with seawater. San Francisco Bay is an example of a tectonic estuary (Figure 4.6.4).
Estuaries are also classified based on their salinity and mixing patterns. The amount of mixing of fresh and salt water in an estuary depends on the rate at which fresh water enters the head of the estuary from river input, and the amount of seawater that enters the estuary mouth as a result of tidal movements. The input of fresh water is reflected in the flushing time of the estuary. This refers to the time it would take for the in-flowing fresh water to completely replace all the fresh water currently in the estuary. Seawater input is measured by the tidal volume, or tidal prism, which is the average volume of sea water entering and leaving the estuary during each tidal cycle. In other words, it is the volume difference between high and low tides. The interaction between the flushing time, tidal volume, and the shape of the estuary will determine the extent and type of water mixing within the estuary.
In a vertically mixed, or well-mixed estuary there is complete mixing of fresh and salt water from the surface to the bottom. In a particular location the salinity is constant at all depths, but across the estuary the salinity is lowest at the head where the fresh water enters, and is highest at the mouth, where the seawater comes in. This type of salinity profile usually occurs in shallower estuaries, where the shallow depths allow complete mixing from the surface to the bottom.
Slightly stratified or partially mixed estuaries have similar salinity profiles to vertically mixed estuaries, where salinity increases from the head to the mouth, but there is also a slight increase in salinity with depth at any point. This usually occurs in deeper estuaries than those that are well-mixed, where waves and currents mix the surface water, but the mixing may not extend all the way to the bottom.
A salt wedge estuary occurs where the outflow of fresh water is strong enough to prevent the denser ocean water to enter through the surface, and where the estuary is deep enough that surface waves and turbulence have little mixing effect on the deeper water. Fresh water flows out along surface, salt water flows in at depth, creating a wedge shaped lens of seawater moving along the bottom. The surface water may remain mostly fresh throughout the estuary if there is no mixing, or it can become brackish depending on the level of mixing that occurs.
Highly stratified profiles are found in very deep estuaries, such as in fjords. Because of the depth, mixing of fresh and salt water only occurs near the surface, so in the upper layers salinity increases from the head to the mouth, but the deeper water is of standard ocean salinity.
Estuaries are very important commercially, as they are home to the majority of the world’s metropolitan areas, they serve as ports for industrial activity, and a large percentage of the world's population lives near estuaries. Estuaries are also very important biologically, especially in their role as the breeding grounds for many species of fish, birds, and invertebrates.
By Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
The most obvious feature of the oceans is that they contain water. Water is so ubiquitous that it may not seem like a very interesting substance, but it has many unique properties that impact global oceanographic and climatological processes. Many of these processes are due to hydrogen bonds forming between water molecules.
The water molecule consists of two hydrogen atoms and one oxygen atom. The electrons responsible for the bonds between the atoms are not distributed equally throughout the molecule, so that the hydrogen ends of water molecules have a slight positive charge, and the oxygen end has a slight negative charge, making water a polar molecule. The negative oxygen side of the molecule forms an attraction to the positive hydrogen end of a neighboring molecule. This rather weak force of attraction is called a hydrogen bond (Figure 5.1.1). If not for hydrogen bonds, water would vaporize at -68o C, meaning liquid water (and thus life) could not exist on Earth. These hydrogen bonds are responsible for some of water’s unique properties:
1. Water is the only substance to naturally exist in a solid, liquid, and gaseous form under the normal range of temperatures and pressures found on Earth. This is due to water’s relatively high freezing and vaporizing points (see below).
2. Water has a high heat capacity, which is the amount of heat that must be added to raise its temperature. Specific heat is the heat required to raise the temperature of 1 g of a substance by 1o C. Water has the highest specific heat of any liquid except ammonia (Table 5.1.1).
Table 5.1.1 Specific heat values for a number of common substances
| Specific Heat (calories/g/Co) | |
|---|---|
| Ammonia | 1.13 |
| Water | 1.00 |
| Acetone | 0.51 |
| Grain Alcohol | 0.23 |
| Aluminum | 0.22 |
| Copper | 0.09 |
| Silver | 0.06 |
Water is therefore one of the most difficult liquids to heat or cool; it can absorb large amounts of heat without increasing its temperature. Remember that temperature reflects the average kinetic energy of the molecules within a substance; the more vigorous the motion, the higher the temperature. In water, the molecules are held together by hydrogen bonds, and these bonds must be overcome to allow the molecules to move freely. When heat is added to water the energy must first go to breaking the hydrogen bonds before the temperature can begin to rise. Therefore, much of the added heat is absorbed by breaking H bonds, not by increasing the temperature, giving water a high heat capacity.
Hydrogen bonds also give water a high latent heat; the heat required to undergo a phase change from solid to liquid, or liquid to gas. The latent heat of fusion is the heat required to go from solid to liquid; 80 cal/g in the case of ice melting to water. Ice is a solid because hydrogen bonds hold the water molecules into a solid crystal lattice (see below). As ice is heated, the temperature rises up to 0o C. At that point, any additional heat goes to melting the ice by breaking the hydrogen bonds, not to increasing the temperature. So as long as ice is present, the water temperature will not increase. This is why your drink will remain cold as long as it contains ice; any heat absorbed goes to melting the ice, not to warming the drink.
When all of the ice is melted, additional heat will increase the temperature of the water 1o C for each calorie of heat added, until it reaches 100o C. At that point, any additional heat goes to overcoming the hydrogen bonds and turning the liquid water into water vapor, rather than increasing the water temperature. The heat required to evaporate liquid water into water vapor is the latent heat of vaporization which has a value of 540 cal/g (Figure 5.1.2).
The high heat capacity of water helps regulate global climate, as the oceans slowly absorb and release heat, preventing rapid swings in temperature (see section 8.1). It also means that aquatic organisms aren't as subjected to the same rapid temperature changes as terrestrial organisms. A deep ocean organism may not experience more than a 0.5o C change in temperature over its entire life, while a terrestrial species may encounter changes of more than 20o C in a single day!
3. Water dissolves more substances than any other liquid; it is a "universal solvent", which is why so many substances are dissolved in the ocean. Water is especially good at dissolving ionic salts; molecules made from oppositely charged ions such as NaCl (Na+ and Cl-). In water, the charged ions attract the polar water molecules. The ions get surrounded by a layer of water molecules, weakening the bond between the ions by up to 80 times. With the bonds weakened between ions, the substance dissolves (Figure 5.1.3).
4. The solid phase is less dense than the liquid phase. In other words, ice floats. Most substances are denser in the solid form than in the liquid form, as their molecules are more closely packed together as a solid. Water is an exception: the density of fresh water is 1.0 g/cm3, while the density of ice is 0.92 g/cm3, and once again, this is due to the action of hydrogen bonds.
As water temperature cools the molecules slow down, eventually slowing enough that hydrogen bonds can form and hold the water molecules in a crystal lattice. The molecules in the lattice are spaced farther apart than the molecules in liquid water, which makes ice less dense than liquid water (Figure 5.1.4). This is familiar to anyone who has ever left a full water bottle in the freezer, only to have it burst as the water freezes and expands.
But the relationship between temperature and water density is not a simple linear one. As water cools, its density increases as expected, as the water molecules slow down and get closer together. However, fresh water reaches its maximum density at a temperature of 4o C, and as it cools beyond that point its density declines as the hydrogen bonds begin to form and the intermolecular spacing increases (Figure 5.1.5 inset). The density continues to decline until the temperature reaches 0o C and ice crystals form, reducing the density dramatically (Figure 5.1.5).
There are a number of important implications to ice being less dense than water. Ice floating on the surface of the ocean helps regulate ocean temperatures, and therefore global climate, by influencing the amount of sunlight that is reflected rather than absorbed (see section 5.6). On a smaller scale, surface ice can prevent lakes and ponds from freezing solid during the winter. As fresh surface water cools, the water gets denser, and sinks to the bottom. The new surface water then cools and sinks, and the process is repeated in what is referred to as overturning, with denser water sinking and less dense water moving to the surface only to be cooled and sink itself. In this way, the entire body of water is cooled somewhat evenly. This process continues until the surface water cools below 4o C. Below 4o C, the water becomes less dense as it cools, so it no longer sinks. Instead, it remains as the surface, getting colder and less dense, until it freezes at 0o C. Once fresh water freezes, the ice floats and insulates the rest of the water beneath it, reducing further cooling. The densest bottom water is still at 4o C, so it does not freeze, allowing the bottom of a lake or pond to remain unfrozen (which is good news for the animals living there) no matter how cold it gets outside.
The dissolved salts in seawater inhibit the formation of the crystal lattice, and therefore make it harder for ice to form. So seawater has a freezing point of about -2o C (depending on salinity), and freezes before a temperature of maximum density is reached. Thus seawater will continue to sink as it gets colder, until it finally freezes.
5. Water has a very high surface tension, the highest of any liquid except mercury (Table 5.1.2). Water molecules are attracted to each other by hydrogen bonds. For molecules not at the water surface, they are surrounded by other water molecules in all directions, so the attractive forces are evenly distributed in all directions. But for molecules at the surface there are few adjacent molecules above them, only below, so all of the attractive forces are directed inwards, away from the surface (Figure 5.1.6). This inwards force is what causes water droplets to take on a spherical shape, and water to bead up on a surface, as the spherical shape provides the minimum possible surface area. These attractive forces also cause the surface of the water to act like an elastic "skin" which allows things like insects to sit on the water's surface without sinking.
Table 5.1.2 Surface tensions of various liquids
| Liquid | Surface Tension (millinewton/meter) | Temperature oC |
|---|---|---|
| Mercury | 487.00 | 15 |
| Water | 71.97 | 25 |
| Glycerol | 63.00 | 20 |
| Acetone | 23.70 | 20 |
| Ethanol | 22.27 | 20 |
By Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
Large waves crashing onto a shore bring a tremendous amount of energy that has a significant eroding effect, and several unique erosion features commonly form on rocky shores with strong waves.
When waves approach an irregular shore, they are slowed down to varying degrees, depending on differences in the water depth, and as they slow, they are bent or refracted (section 3.3). In Figure 4.3.1, wave energy is represented by the blue arrows. That energy is evenly spaced out in the deep water, but because of refraction, the energy of the waves is being focused on the headlands. On irregular coasts, the headlands receive much more wave energy than the intervening bays, and thus they are more strongly eroded. The result of this is coastal straightening, where an irregular coast will eventually become straightened, although that process may take millions of years.
Wave erosion is greatest in the surf zone, where the wave base is impinging strongly on the seafloor and where the waves are breaking. The result is that the substrate in the surf zone is typically eroded to a flat surface known as a wave-cut platform (or wave-cut terrace) (Figure 4.3.2). A wave-cut platform extends across the intertidal zone.
Arches and sea caves form as a result of the erosion of relatively non-resistant rock. Wave action and strong longshore currents can carve a cave into a headland, and if the erosion extends all the way through, it becomes an arch. If a hole develops in the ceiling of a cave, a blowhole can be created, shooting water into the air when waves crash in the cave. An arch in the Barachois River area of western Newfoundland, Canada, is shown in Figure 4.3.3. This feature started out as a sea cave, and then, after being eroded from both sides, became an arch. During the winter of 2012-2013, the arch collapsed, leaving a small stack at the end of the point.
The tower of rock left behind from a collapsed arch is called a sea stack (Figure 4.3.4). But sea stacks can also form during the formation of wave-cut platforms or other features, when relatively resistant rock that does not get completely eroded remains behind to form the stack.
Our modern understanding of tide formation stems from Isaac Newton's Law of Universal Gravitation, which states that any two objects have a gravitational attraction to each other. The magnitude of the force is proportional to the masses of the objects, and inversely proportional to the square of the distance between the objects, according to the equation in Figure 3.5.1.
In the case of tides, there are a few other factors that modify this equation so that the distance (r) is cubed rather than squared, giving distance an even greater impact on tidal forces. But for our purposes, the important lesson is that the greater the masses of the objects, the greater the gravitational force, and the farther the objects are from each other, the weaker the force.
Such a gravitational force exists between the Earth and moon, attempting to pull them towards each other. Since the water covering Earth is fluid (unlike the solid land that is more resistant to tidal forces), this gravitational force pulls water towards the moon, creating a "bulge" of water on the side of the Earth facing the moon (Figure 3.5.2). This bulge always faces the moon, while the Earth rotates through it; the regions of Earth moving through the bulge experience a high tide, while those parts of the Earth away from the bulge experience a low tide.
If the tides were this simple, everywhere on Earth would see one high tide per day, as there would only be a bulge of water on the side closest to the moon. However, if you have ever looked at tide charts, or lived near the ocean, you probably know that in most places there are two high tides and two low tides per day. Where is this second high tide "bulge" coming from?
The gravitational force between the Earth and moon might be expected to draw the two objects closer together, however, this is not happening. This is because the inward gravitational force is opposed by outward forces that keep the Earth and moon apart. The outward force is an intertial force created by the rotation of the Earth and moon. Contrary to popular belief, the moon is not simply rotating around the Earth; in fact, the Earth and moon are both rotating around each other. Imagine the Earth and moon as equal-sized objects revolving around a point at their center of mass. If both objects had the same mass, the center of rotation would be a point equidistant between the two objects. But since the mass of the Earth is 82 times greater than the mass of the moon, the center of revolution must be closer to the Earth. As an analogy, think about two people on a see-saw. If the people are of roughly equal size, they can sit on either end of the see-saw at it will rotate around a point at equal distance between them. But if the two people have very different masses, such as a large adult and a small child, the larger person must move closer to the pivot point for the see-saw to rotate effectively. In the same way, the center of rotation between the Earth and the moon (the barycenter) must be located closer to the Earth. In fact, the center of rotation lies within the Earth, about 1600 km below the surface. As the Earth and moon rotate around the barycenter, the moon travels much farther than the Earth, giving the impression that the moon is rotating around Earth (Figure 3.5.3).
The rotation of the Earth-moon system creates an outward inertial force, which balances the gravitational force to keep the two bodies in their orbits. The inertial force has the same magnitude everywhere on Earth, and is always directed away from the moon. Gravitational force, on the other hand, is always directed towards the moon, and is stronger on the side of the Earth closest to the moon. Figure 3.5.4 describes how these forces combine to create the tidal forces. At point O in the center of the Earth, the gravitational force (Fg) and the inertial force (Fr) are equal, and cancel each other out. On the side of Earth closest to the moon, the inward gravitational force (Fg) is greater than the outward inertial force (Fr); the net resulting force (A) is directed towards the moon, and creates a bulge of water on the side facing the moon. On the side of Earth opposite the moon, the outward inertial force is greater than the inward gravitational force; the net resulting force (C) is directed away from the moon, creating a water bulge directed away from the moon.
Now, as the Earth rotates through a 24 hour day, each region passes through two bulges, and experiences two high tides and two low tides per day. This represents Newton's Equilibrium Theory of Tides, where there are two high tides and two low tides per day, of similar heights, each six hours apart. But as with everything else in oceanography, reality is much more complex than this idealized situation.
Some of the additional complexity is because in addition to the moon, the sun also exerts tide-affecting forces on Earth. The solar gravitational and inertial forces arise for the same reasons described above for the moon, but the magnitudes of the forces are different. The sun is 27 million times more massive than the moon, but it is 387 times farther away from the Earth. Despite its larger mass, because the sun is so much farther away than the moon, the sun's gravitational forces are only about half as strong as the moon's (remember that distance is cubed in the gravity equation). The sun thus creates its own, smaller water bulges, independent of the moon's, that contribute to the creation of tides.
When the sun, Earth and moon are aligned, as occurs during new and full moons, the solar and lunar bulges are also aligned, and add to each other (constructive interference; see section 4.2) creating an especially high tidal range; high high tides and low low tides (Figure 3.5.5). This period of maximum tidal range is called a spring tide, and they occur every two weeks.
When the sun, Earth and moon are at 90o to each other, the solar and lunar bulges are out of phase, and cancel each other out (destructive interference). Now the tidal range is small, with low high tides and high low tides (Figure 3.5.6). These are neap tides, and occur every two weeks, when the moon is in its 1/4 and 3/4 phases (Figure 3.5.7).
By Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
We learned in section 3.3 that refraction causes waves to approach parallel to shore. However, most waves still reach the shore at a small angle, and as each one arrives, it pushes water along the shore, creating what is known as a longshore current within the surf zone (the areas where waves are breaking) (Figure 4.2.1). Longshore currents can move up to 4 km/hr, strong enough to carry people with them, as everyone knows who has been swimming in the ocean only to look up and see that they have been carried far down the beach from their towel!
Learning Objectives
After reading this chapter you should be able to:
- identify the parts of a basic wave
- define the terminology used to describe the motion of a wave (i.e. period, frequency, speed etc.)
- explain the circular motion of water particles involved in wave motion
- explain the difference between deep water waves and shallow water waves
- identify the factors that influence wave speed in deep and shallow waves
- identify the factors that determine the energy of wind-generated waves
- define the concept of restoring force
- define significant wave height
- explain the creation of ocean swell
- define the concepts of destructive, constructive and mixed interference
- explain why waves break as they approach shore
- explain the differences in the different types of breakers, and how the bottom topography impacts breaker type
- explain why waves approach parallel to shore, and why waves are larger off of points and smaller in bays
- explain how tsunamis are formed, and how they behave in the ocean
- explain Newton's Law of Universal Gravitation and how it applies to tides
- explain why most places on Earth experience two tides per day, not just the one predicted from gravitational attraction between the Earth and moon
- explain how the Earth, sun and moon interact to create spring and neap tides
- explain why the sun has a smaller effect on tides than the moon
- explain why tides do not occur at the same time every day
- explain the concept of amphidromic circulation
- identify diurnal, semi-diurnal, and mixed tides
- identify the phases of a tidal current
- define a tidal bore
Waves come in many shapes and sizes; a 100 foot wave might be a surfer's dream, but a ship captain's nightmare. What was the largest wave ever recorded? 50 feet? 100 feet? Not even close. That record belongs to a wave created in Lituya Bay, Alaska, on July 9, 1958 (Figure 3.1). On that day, a magnitude 7.8 earthquake caused a massive rockslide that slid down a mountainside and into the headwaters of the bay. The rockslide created a splash wave that was high enough to flatten vegetation up to 1722 ft (525 m) above sea level! The wave then moved through the narrow bay towards the sea, destroying a number of fishing boats along the way. Miraculously, a father and son on one fishing boat were carried above the trees by the wave, and survived to tell the story. This is by far the largest wave, a megatsunami, ever reliably recorded. The waves we will discuss in this chapter may not be quite that dramatic, but it is still important to know how they form, how they are propagated, and what happens to them as they interact with the shore.
After discussing various types of waves at sea and along the shore, we will discuss tides. However, at least in terms of wavelength, the largest waves in the ocean are the tides, where one wavelength stretches halfway around the Earth. The crests of these long waves represent the high tides, while the troughs create low tides.
You probably learned when you were younger that the basic cause of the tides is the gravitational attraction between the Earth and moon. This is a very old idea, as the Greek scientist Pytheas first made the connection between the tides and the moon back in 330 B.C.E. Isaac Newton's gravitational work the 1600s led to our modern understanding of tidal cycles, however, we now know that the tides involve a lot more than just the Earth and the moon. There are many variables that influence the tides, yet despite this complexity, we are able to create accurate tide charts predicting the heights and timing of tides months or even years in advance.
Modified from Paul Webb, used under a CC-BY 4.0 international license. Download this book for free at https://rwu.pressbooks.pub/webboceanography/front-matter/preface/
Divergent boundaries are spreading boundaries, where new oceanic crust is created to fill in the space as the plates move apart. Most divergent boundaries are located along mid-ocean oceanic ridges (although some are on land). The mid-ocean ridge system is a giant undersea mountain range, and is the largest geological feature on Earth; at 65,000 km long and about 1000 km wide, it covers 23% of Earth’s surface (Figure 2.5.1). Because the new crust formed at the plate boundary is warmer than the surrounding crust, it has a lower density so it sits higher on the mantle, creating the mountain chain. Running down the middle of the mid-ocean ridge is a rift valley 25-50 km wide and 1 km deep. Although oceanic spreading ridges appear to be curved features on Earth’s surface, in fact the ridges are composed of a series of straight-line segments, offset at intervals by faults perpendicular to the ridge, called transform faults. These transform faults make the mid-ocean ridge system look like a giant zipper on the seafloor (Figure 2.5.2). As we will see in section 3.7, movements along transform faults between two adjacent ridge segments are responsible for many earthquakes.
The crustal material created at a spreading boundary is always oceanic in character; in other words, it is igneous rock (e.g., basalt or gabbro, rich in ferromagnesian minerals), forming from magma derived from partial melting of the mantle caused by decompression as hot mantle rock from depth is moved toward the surface (Figure 2.5.3). The triangular zone of partial melting near the ridge crest is approximately 60 km thick and the proportion of magma is about 10% of the rock volume, thus producing crust that is about 6 km thick. This magma oozes out onto the seafloor to form pillow basalts, breccias (fragmented basaltic rock), and flows, interbedded in some cases with limestone or chert. Over time, the igneous rock of the oceanic crust gets covered with layers of sediment, which eventually become sedimentary rock.
Spreading is hypothesized to start within a continental area with up-warping or doming of crust related to an underlying mantle plume or series of mantle plumes. The buoyancy of the mantle plume material creates a dome within the crust, causing it to fracture. When a series of mantle plumes exists beneath a large continent, the resulting rifts may align and lead to the formation of a rift valley (such as the present-day Great Rift Valley in eastern Africa). It is suggested that this type of valley eventually develops into a linear sea (such as the present-day Red Sea), and finally into an ocean (such as the Atlantic). It is likely that as many as 20 mantle plumes, many of which still exist, were responsible for the initiation of the rifting of Pangaea along what is now the mid-Atlantic ridge.
There are multiple lines of evidence demonstrating that new oceanic crust is forming at these seafloor spreading centers:
1. Age of the crust:
Comparing the ages of the oceanic crust near a mid-ocean ridge shows that the crust is youngest right at the spreading center, and gets progressively older as you move away from the divergent boundary in either direction, aging approximately 1 million years for every 20-40 km from the ridge. Furthermore, the pattern of crust age is fairly symmetrical on either side of the ridge (Figure 2.5.4).
The oldest oceanic crust is around 280 Ma in the eastern Mediterranean, and the oldest parts of the open ocean are around 180 Ma on either side of the north Atlantic. It may be surprising, considering that parts of the continental crust are close to 4,000 Ma old, that the oldest seafloor is less than 300 Ma. Of course, the reason for this is that all seafloor older than that has been either subducted (see section 3.6) or pushed up to become part of the continental crust. As one would expect, the oceanic crust is very young near the spreading ridges (Figure 2.5.4), and there are obvious differences in the rate of sea-floor spreading along different ridges. The ridges in the Pacific and southeastern Indian Oceans have wide age bands, indicating rapid spreading (approaching 10 cm/year on each side in some areas), while those in the Atlantic and western Indian Oceans are spreading much more slowly (less than 2 cm/year on each side in some areas).
2. Sediment thickness:
With the development of seismic reflection sounding (similar to echo sounding described in section 1.4) it became possible to see through the seafloor sediments and map the bedrock topography and crustal thickness. Hence sediment thicknesses could be mapped, and it was soon discovered that although the sediments were up to several thousands of meters thick near the continents, they were relatively thin — or even non-existent — in the ocean ridge areas (Figure 2.5.5). This makes sense when combined with the data on the age of the oceanic crust; the farther from the spreading center the older the crust, the longer it has had to accumulate sediment, and the thicker the sediment layer. Additionally, the bottom layers of sediment are older the farther you get from the ridge, indicating that they were deposited on the crust long ago when the crust was first formed at the ridge.
3. Heat flow:
Measurements of rates of heat flow through the ocean floor revealed that the rates are higher than average (about 8x higher) along the ridges, and lower than average in the trench areas (about 1/20th of the average). The areas of high heat flow are correlated with upward convection of hot mantle material as new crust is formed, and the areas of low heat flow are correlated with downward convection at subduction zones.
4. Magnetic reversals:
In section 2.2 we saw that rocks could retain magnetic information that they acquired when they were formed. However, Earth's magnetic field is not stable over geological time. For reasons that are not completely understood, the magnetic field decays periodically and then becomes re-established. When it does re-establish, it may be oriented the way it was before the decay, or it may be oriented with the reversed polarity. During periods of reversed polarity, a compass would point south instead of north. Over the past 250 Ma, there have a few hundred magnetic field reversals, and their timing has been anything but regular. The shortest ones that geologists have been able to define lasted only a few thousand years, and the longest one was more than 30 million years, during the Cretaceous (Figure 2.5.6). The present “normal” event has persisted for about 780,000 years.
Beginning in the 1950s, scientists started using magnetometer readings when studying ocean floor topography. The first comprehensive magnetic data set was compiled in 1958 for an area off the coast of British Columbia and Washington State. This survey revealed a mysterious pattern of alternating stripes of low and high magnetic intensity in sea-floor rocks (Figure 2.5.7). Subsequent studies elsewhere in the ocean also observed these magnetic anomalies, and most importantly, the fact that the magnetic patterns are symmetrical with respect to ocean ridges. In the 1960s, in what would become known as the Vine-Matthews-Morley (VMM) hypothesis, it was proposed that the patterns associated with ridges were related to the magnetic reversals, and that oceanic crust created from cooling basalt during a normal event would have polarity aligned with the present magnetic field, and thus would produce a positive anomaly (a black stripe on the sea-floor magnetic map), whereas oceanic crust created during a reversed event would have polarity opposite to the present field and thus would produce a negative magnetic anomaly (a white stripe). The widths of the anomalies varied according to the spreading rates characteristic of the different ridges. This process is illustrated in Figure 2.5.8. New crust is formed (panel a) and takes on the existing normal magnetic polarity. Over time, as the plates continue to diverge, the magnetic polarity reverses, and new crust formed at the ridge now takes on the reversed polarity (white stripes in Figure 2.5.8). In panel b, the poles have reverted to normal, so once again the new crust shows normal polarity before moving away from the ridge. Eventually, this creates a series of parallel, alternating bands of reversals, symmetrical around the spreading center (panel c).
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- a streamlined body shape for more efficient movement through the water
- a high metabolic rate that generates a lot of heat and layers of fat and, in some cases, fur, to conserve this heat
- modifications to their respiratory system to collect and retain large volumes of oxygen to allow deep and repetitive dives
- osmotic adaptations that free them from any requirement for fresh water
Order Cetacea: the whales
- Whales also have special adaptations, such as:
- A thick layer of blubber
- Their nostrils have migrated to the top of the head so that the animal doesn't have to come completely out of the water to breathe
- Large, deeply convoluted brains
- A special skin that "gives" to dampen out irregularities in water flow
- Bronchial cartilage that supports their lungs against pressure during deep dives
- Their blood is especially rich in hemoglobin to store more oxygen for diving
- They can slow their heartbeat while diving
- The blood supply can be reduced to all but the vital organs while diving, to conserve oxygen
Whale groups come in two 'flavors' (sub-Orders) – Odontocetes (toothed whales) and Mysticetes (baleen whales)
- Sperm whales are the largest of the toothed whales, reaching lengths up to 60 ft
- It also possesses the largest brain ever to have evolved on Earth
- They can dive to greater depths than any other air-breathing animal – to 3000 ft – and can stay down for over an hour
- Their head is filled with up to a ton of clear oil, presumably used to focus sound waves passing to their prey and back
Mysticetes (baleen whales): Also known as baleen whales, the Mysticetes are considered to be more highly evolved than the Odontocetes. Baleen is a bristly, bush, fibrous substance set in the jaws in overlapping plates; it looks like a gigantic comb.
- They feed on small planktonic arthropods (krill)
- Baleen is a straining mechanism
- The whale takes in a big gulp of water, closes its jaws, raises its tongue to expel the water, and the food is caught on the surface of the baleen
- The food is swallowed after being licked from the baleen by the whale's giant tongue
- Baleen whales include the blue, humpback and gray, among others
- The blue whale is probably the largest animal ever to have lived on Earth
- They're usually described as up to about 100 feet in length and up to about 200 tons
- this compares to three railroad cars in length and about 1600 people in weight
- Blue whales feed on krill, which are tiny shrimp-like crustaceans
- A mature blue whale consumes about 4 tons of krill per day when it's feeding
- For a 7 to 8 month period; it fasts when it is in tropical waters
- From December through April, the blue whale gorges on krill around the Antarctic
Humpback whale
- Humpbacks are stocky and seldom exceed 50 ft in length
- They're generally black above and white below, with extremely long, wing-like flippers
- They've become known for their "song", which are sounds with definite patterns and sequences
- All the whales in a given population sing the same song, but the song changes every year
- Humpbacks are also very acrobatic, at times leaping completely out of the water
- It is estimated that there are no more than 10,000 humpbacks surviving today
Gray whale
- The Gray whale is the only large whale with a heavily mottled appearance and a knobby ridge down the back
- Its 'natural' color is dark gray or black, but it's covered by a profusion of spots, scars, patches, and clusters of barnacles, which gives it a mottled, grayish appearance
- The Gray whale makes the longest migration of any mammal, an annual round-trip of some 10,000 miles from the Bering and Chukchi seas in the high Arctic to the warm lagoons of Baja California
- They were easily caught by whalers and almost became extinct in the 1940's
- Given full protection in 1946, they've made a successful comeback and their population, estimated at about 21,000, is now believed to be inline with the carrying capacity of their range
- The Gray whale is the most heavily parasitized of all cetaceans – playing host to three species of lice, some over an inch long – in one case, more than 100,000 of one species of louse were removed from a single whale
Dr. Cristina Cardona
Continental margins refer to the region of transition from the land to the deep seafloor, i.e. between continental and oceanic crust. In an active continental margin, the boundary between the continent and the ocean is also a tectonic plate boundary, so there is a lot of geological activity around the margin. The west coast of the United States is an example of an active margin, where the coastline corresponds with the boundary between the Pacific and North America Plates. A passive continental margin occurs where the transition from land to sea is not associated with a plate boundary. The east coast of the United States is a good example; the plate boundary is located along the mid Atlantic ridge, far from the coast. Passive margins are less geologically active. Figure 1.2.1 shows an idealized passive margin. When examining this figure, and others like it, note that there is significant vertical exaggeration; the depth scale covers approximately 5000 m, while the horizontal scale extends around 300 km. This makes the features look much steeper than they actually are. The bar at the bottom of Figure 1.2.1 shows what a passive margin would look like without this exaggeration; there is a much more gradual transition to depth.
The continental shelf is the shallow, flooded edge of the continent. Geologically the shelf is still part of the continental crust, but it is often overlaid with marine sediments. On average, the shelf extends about 80 km from the coast; some margins have very little shelf, while the Siberian Shelf in the Arctic extends roughly 1500 km. The depth of the shelf generally remains below about 150 m, and the floor of the shelf is fairly flat. The flat topography is the result of changes in sea level; throughout history the shelves have been both submerged and exposed, and as sea level rose and fell, wave action, ice sheets, and other erosional processes smoothed out the shelf surface. Wave action and the movement of sediments over the shelf have continued this smoothing process. Continental shelves only make up about 6% of the ocean's surface area, but they are biologically one of the richest parts of the ocean; their shallow depth prevents nutrients from sinking out, and their proximity to the coast provides significant nutrient input. The continental shelf ends at the shelf break, which is the point where the angle of the seafloor begins to get steeper. The shelf break averages about 135 m deep.
After the shelf break, the seafloor takes on a steeper angle (about 4o) as it descends to the deep ocean. This steeper portion of the margin is the continental slope, and it extends from the shelf break down to 3000-5000m. In some parts of the ocean, large submarine canyons have been carved into the continental slope; for example, Monterey Canyon in Monterey Bay, California, is a submarine canyon similar in size to the Grand Canyon! These canyons may be carved out by turbidity currents, which are essentially landslides of sediment, rocks, and other debris down the face of the slope.
At the bottom of the slope is the continental rise. This area represents where the continental crust meets the oceanic crust, as the slope begins to level off to become the deep ocean floor. The rise consists of a thick layer of accumulated sediment coming from the continent, so it is difficult to tell where the slope ends and the rise begins.
After the rise comes the abyssal plain, or the deep ocean floor, lying between 4500 - 6000 m. The abyssal plain includes most of the ocean floor, and is the flattest region on Earth. It is flat due to millions of years of sediment accumulation on the bottom, which buries many bottom features (Figure 1.2.2).
Passive margins, as described above, have wide shelves, gentle slopes, and a well-developed rise. Since passive margins are not plate boundaries, they experience long periods of relative stability which can lead to the development of these features. Active margins have similar features to passive margins, but the plate boundary affects the properties of the features. Active margins, like the Pacific coast of North America, have narrower shelves, steeper slopes, and little to no rise, particularly in convergent boundaries. Trenches associated with subduction zones act as sediment traps, preventing the accumulation of a continental rise, and keeping sediments off of the abyssal plains.
Large waves crashing onto a shore bring a tremendous amount of energy that has a significant eroding effect, and several unique erosion features commonly form on rocky shores with strong waves.
When waves approach an irregular shore, they are slowed down to varying degrees, depending on differences in the water depth, and as they slow, they are bent or refracted (section 3.3). In Figure 4.3.1, wave energy is represented by the blue arrows. That energy is evenly spaced out in the deep water, but because of refraction, the energy of the waves is being focused on the headlands. On irregular coasts, the headlands receive much more wave energy than the intervening bays, and thus they are more strongly eroded. The result of this is coastal straightening, where an irregular coast will eventually become straightened, although that process may take millions of years.
Wave erosion is greatest in the surf zone, where the wave base is impinging strongly on the seafloor and where the waves are breaking. The result is that the substrate in the surf zone is typically eroded to a flat surface known as a wave-cut platform (or wave-cut terrace) (Figure 4.3.2). A wave-cut platform extends across the intertidal zone.
Arches and sea caves form as a result of the erosion of relatively non-resistant rock. Wave action and strong longshore currents can carve a cave into a headland, and if the erosion extends all the way through, it becomes an arch. If a hole develops in the ceiling of a cave, a blowhole can be created, shooting water into the air when waves crash in the cave. An arch in the Barachois River area of western Newfoundland, Canada, is shown in Figure 4.3.3. This feature started out as a sea cave, and then, after being eroded from both sides, became an arch. During the winter of 2012-2013, the arch collapsed, leaving a small stack at the end of the point.
The tower of rock left behind from a collapsed arch is called a sea stack (Figure 4.3.4). But sea stacks can also form during the formation of wave-cut platforms or other features, when relatively resistant rock that does not get completely eroded remains behind to form the stack.
Generally ocean temperatures range from about -2o to 30o C (28-86o F). The warmest water tends to be surface water in low latitude regions, while the surface water at the poles is obviously much colder (Figure 5.6.1). Note that at equivalent latitudes, water on the eastern side of the ocean basins is colder than the water on the western side. This has to do with the pattern of surface currents. Even though surface water can be quite warm, most of the water in the oceans is deeper, colder water, so that the average temperature of the entire ocean is about 4o C, which is roughly the temperature inside your refrigerator.
A typical temperature profile for open ocean, mid-latitude water is shown in Figure 5.6.2. Water is warmest at the surface, as it is warmed by the sun, and the sun's rays can only penetrate depths less than 1000 m. Since the surface water is warmer it is also less dense than the deep water, so it remains at the surface where it can be warmed even more. Temperature is fairly constant in the upper 100-200 m in what is called the mixed layer. The mixed layer results from surface winds, waves, and currents that mix the upper water and distribute the heat throughout this layer. Below the mixed layer there is a rapid decline in temperature over a fairly narrow increase in depth. This is called the thermocline. Below the thermocline the deep ocean temperature is fairly constant at about 2o C, continuing down to the bottom. There is little temperature change in the deep ocean, as it is far removed from significant heat sources, making it one of the most thermally stable regions on earth. Temperature may fluctuate by less than half a degree per year in the deep ocean (Figure 5.6.3).
Temperature profiles vary at different latitudes, as the surface water is warmer near the equator and colder at the poles. In low latitude tropical regions the sea surface is much warmer, leading to a highly pronounced thermocline (Figure 5.6.4). Additionally, there is not much seasonal change in surface temperature in tropical regions, so there is little seasonal change in the profiles. In high latitude (polar) regions, there is little difference between the surface temperature and the deep water temperature, and temperature is fairly constant (and cold) at all depths. Polar waters therefore lack a strong thermocline, and as with tropical water, there is little seasonal change in temperatures. Mid-latitude temperate regions show greater seasonal fluctuations in surface temperature than the poles or the tropics; an 8-15o C difference from summer to winter in temperate zones, compared to only ~2o C in polar and tropical areas. In temperate regions, the surface water is much warmer in the summer and the thermocline is more pronounced compared to the winter months. But in the winter the thermocline is deeper at mid-latitudes than it is in the summer. This is because winter storms churn up the surface water more than occurs in the summer, creating a deeper mixed layer and thus a deeper thermocline (Figure 5.6.5).
Due to the high heat capacity of water, daily fluctuations in ocean temperature are fairly insignificant.
Convergent boundaries, where two plates are moving toward each other, are of three types, depending on the type of crust present on either side of the boundary — oceanic or continental. The types are ocean-ocean, ocean-continent, and continent-continent.
At an ocean-ocean convergent boundary, one of the plates (oceanic crust and lithospheric mantle) is pushed, or subducted, under the other (Figure 2.6.1). Often it is the older and colder plate that is denser and subducts beneath the younger and warmer plate. There is commonly an ocean trench along the boundary as the crust bends downwards. The subducted lithosphere descends into the hot mantle at a relatively shallow angle close to the subduction zone, but at steeper angles farther down (up to about 45°). The significant volume of water within the subducting material is released as the subducting crust is heated. It mixes with the overlying mantle, and the addition of water to the hot mantle lowers the crust’s melting point and leads to the formation of magma (flux melting). The magma, which is lighter than the surrounding mantle material, rises through the mantle and the overlying oceanic crust to the ocean floor where it creates a chain of volcanic islands known as an island arc. A mature island arc develops into a chain of relatively large islands (such as Japan or Indonesia) as more and more volcanic material is extruded and sedimentary rocks accumulate around the islands. Earthquakes occur relatively deep below the seafloor, where the subducting crust moves against the overriding crust.
Examples of ocean-ocean convergent zones are subduction of the Pacific Plate south of Alaska (creating the Aleutian Islands) and under the Philippine Plate, where it creates the Marianas Trench, the deepest part of the ocean.
At an ocean-continent convergent boundary, the denser oceanic plate is pushed under the less dense continental plate in the same manner as at an ocean-ocean boundary. Sediment that has accumulated on the seafloor is thrust up into an accretionary wedge, and compression leads to thrusting within the continental plate (Figure 2.6.2). The magma produced adjacent to the subduction zone rises to the base of the continental crust and leads to partial melting of the crustal rock. The resulting magma ascends through the crust, producing a mountain chain with many volcanoes. As with an ocean-ocean boundary, the subducting crust can produce a deep trench running parallel to the coastline.
Examples of ocean-continent convergent boundaries are subduction of the Nazca Plate under South America (which has created the Andes Mountains and the Peru Trench) and subduction of the Juan de Fuca Plate under North America (creating the Cascade Range).
A continent-continent collision occurs when a continent or large island that has been moved along with subducting oceanic crust collides with another continent (Figure 2.6.3). The colliding continental material will not be subducted because it is too light (i.e., because it is composed largely of light continental rocks), but the root of the oceanic plate will eventually break off and sink into the mantle. There is tremendous deformation of the pre-existing continental rocks, forcing the material upwards and creating mountains.
Examples of continent-continent convergent boundaries are the collision of the India Plate with the Eurasian Plate, creating the Himalaya Mountains, and the collision of the African Plate with the Eurasian Plate, creating the series of ranges extending from the Alps in Europe to the Zagros Mountains in Iran.
Radiant energy from the sun is important for several major oceanic processes:
- Climate, winds, and major ocean currents are ultimately dependent on solar radiation reaching the Earth and heating different areas to different degrees.
- Sunlight warms the surface water where much oceanic life lives.
- Solar radiation provides light for photosynthesis, which supports the entire ocean ecosystem.
The energy reaching Earth from the sun is a form of electromagnetic radiation, which is represented by the electromagnetic spectrum (Figure 5.9.1). Electromagnetic waves vary in their frequency and wavelength. High frequency waves have very short wavelengths, and are very high energy forms of radiation, such as gamma rays and x-rays. These rays can easily penetrate the bodies of living organisms and interfere with individual atoms and molecules. At the other end of the spectrum are low energy, long wavelength waves such as radio waves, which do not pose a hazard to living organisms.
Most of the solar energy reaching the Earth is in the range of visible light, with wavelengths between about 400-700 nm. Each color of visible light has a unique wavelength, and together they make up white light. The shortest wavelengths are on the violet and ultraviolet end of the spectrum, while the longest wavelengths are at the red and infrared end. In between, the colors of the visible spectrum comprise the familiar "ROYGBIV"; red, orange, yellow, green, blue, indigo, and violet.
Water is very effective at absorbing incoming light, so the amount of light penetrating the ocean declines rapidly (is attenuated) with depth (Figure 5.9.2). At 1 m depth, only 45% of the solar energy that falls on the ocean surface remains. At 10 m depth only 16% of the light is still present, and only 1% of the original light is left at 100 m. No light penetrates beyond 1000 m.
In addition to overall attenuation, the oceans absorb the different wavelengths of light at different rates (Figure 5.9.2). The wavelengths at the extreme ends of the visible spectrum are attenuated faster than those wavelengths in the middle. Longer wavelengths are absorbed first; red is absorbed in the upper 10 m, orange by about 40 m, and yellow disappears before 100 m. Shorter wavelengths penetrate further, with blue and green light reaching the deepest depths.
This explains why everything appears blue under water. The colors we perceive depends on the wavelengths of light that are received by our eyes. If an object appears red to us, that is because the object reflects red light but absorbs all of the other colors. So the only color reaching our eyes is red. Under water, blue is the only color of light still available at depth, so that is the only color that can be reflected back to our eyes, and everything has a blue tinge under water. A red object at depth will not appear red to us because there is no red light available to reflect off of the object. Objects in water will only appear as their real colors near the surface where all wavelengths of light are still available, or if the other wavelengths of light are provided artificially, such as by illuminating the object with a dive light.
Water in the open ocean appears clear and blue because it contains much less particulate matter, such as phytoplankton or other suspended particles, and the clearer the water, the deeper the light penetration. Blue light penetrates deeply and is scattered by the water molecules, while all other colors are absorbed; thus the water appears blue. On the other hand, coastal water often appears greenish (Figure 5.9.2). Coastal water contains much more suspended silt and algae and microscopic organisms than the open ocean. Many of these organisms, such as phytoplankton, absorb light in the blue and red range through their photosynthetic pigments, leaving green as the dominant wavelength of reflected light. Therefore the higher the phytoplankton concentration in water, the greener it appears. Small silt particles may also absorb blue light, further shifting the color of water away from blue when there are high concentrations of suspended particles.
The ocean can be divided into depth layers depending on the amount of light penetration, as discussed in section 1.3 (Figure 5.9.3). The upper 200 m is referred to as the photic or euphotic zone. This represents the region where enough light can penetrate to support photosynthesis, and it corresponds to the epipelagic zone. From 200-1000 m lies the dysphotic zone, or the twilight zone (corresponding with the mesopelagic zone). There is still some light at these depths, but not enough to support photosynthesis. Below 1000 m is the aphotic (or midnight) zone, where no light penetrates. This region includes the majority of the ocean volume, which exists in complete darkness.