Introductory Oceanography (OCNG 251)

Study Guide: Part 1 This half session dealt with the construction of all conditions responsible for the observed global circulation patterns in the World Ocean. In a sense, we started the course from the very end, trying to build an ocean and understanding its physical structure (the water part in this case). The objective of the entire session is to understand the concept behind Figure 1 below:

Figure 1: General circulation pattern of the ocean. Surface currents are indicated in red while deep currents are presented in blue.

In Figure 1, one can see that there is a link between surface circulation (red) and deep circulation (blue). Of course, to conserve mass, there must be a link between these wo circulation patterns. Areas of “deep water formation” will transfer water from the surface to the deep ocean whereas water returns to the surface via zones of upwelling. Transfer from the surface to the deep ocean will occur due to densification (increased density) of surface water (mostly through cooling but also through some increased salinity during ice formation and salt concentration in seawater). Upwelling will occur through physical transfer from current formation (Ekman circulation in eastern ocean basins) and as water is pushed up continental slope (like when the North Atlantic Deep Water is pushed up the slope of the Antarctic continent). These features are all shown in Figure 1 with areas of deep water formation as purple dots (North and South Atlantic), and areas of upwelling with blue‐to‐red arrows (eastern regions of ocean basins). Also note that surface currents are characterized by circular patterns, called gyres, in each oceanic basin for each hemisphere (Atlantic and Pacific each have 2 gyres, whereas the Indian has only 1).

The entire purpose of the first session was thus to bring all the elements necessary to comprehend the processes responsible for the ocean circulation illustrated in Figure 1. These elements are: OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 2

‐ Systems and cycles. Specifically, how mass and energy cycle through different section of a system (from the micro‐ to macro‐scales). In this section we emphasized notions of reservoir, flux, source/sink, residence time, steady state, as well as positive and negative feedback mechanisms. ‐ Physical properties of water and, in particular, how temperature and salinity affect the density of seawater. We also focused on heat capacity to explain the temperature changes different media experience (i.e. atmosphere vs. ocean, continents, vs. oceans, etc) when subjected to a gain or loss of heat. ‐ Heat budget of the earth, particularly with respect to the unbalance in incoming short wave radiations and outgoing long wave radiations that is observed in inter‐tropical vs. high latitude zones. ‐ Atmospheric circulation, as it is driven by that same unbalance in the earth heat budget and affected by the earth’s rotation ( Coriolis). The interplay of these processes then leads to global as well as seasonal wind patterns (e.g. easterlies/westerlies and monsoons, respectively). ‐ Surface ocean circulation, driven itself by the wind drag of constant winds and affected by coriolis, vorticity, and geostrophic forces. Except for the effect of local winds, the general surface ocean circulation follows the atmospheric High/Low distribution pattern with circular motion (gyres) in each ocean basin. The circulation is clockwise in the north hemisphere, and counterclockwise in the south hemisphere. ‐ Deep ocean circulation, driven by density formation in high latitude zones. Surface water can undergo large increases in density due to an interplay of salinity and temperature changes. When warm water cools, its density increases markedly. Similarly, when water increases in salinity, its density increases as well. The cooling of surface seawater in northern latitudes (e.g. sub‐Arctic seas and around Antartica) leads to an increase in its density and thus vertical transfer of water towards the deep ocean. Similarly, during sea ice formation, the expulsion of salts from the forming ice results in brine formation (increase in salinity in sea waters) and thus an increase in the water density. These processes lead to deep water‐mass formation, each with specific density conditions that help or prevent their mixing in the deep ocean. ‐ Global ocean circulation. The surface and deep ocean circulations are tied at both “ends” where surface water cools (at high latitudes) to form deep waters, and where deep waters are upwelled towards to surface (mostly on eastern boundaries of oceans) to reintegrate the surface circulation loops and eventually reach the cooling sites for another cycle. On average, a full ocean circulation cycle takes several hundred years to complete (~500 yrs) but this “mixing speed” is variable and can accelerate or decelerate depending on the rate of deep water formation (cooling, salinity changes) and upwelling (wind strength, atmospheric pressure oscillation). ‐ Earth climate balance. The relationship between atmospheric and ocean circulation, help redistribute heat from zones of surplus radiation (inter‐tropical zones) to zones of deficit (high latitudes). In low latitudes, the majority of the heat transfer occurs through ocean circulation, whereas atmospheric circulation is responsible for most of the heat transfer in mid‐ to high latitudes. Event such as hurricanes are rapid and natural “pressure valve” processes that transfer large amounts of heat from the inter‐tropical zones to mid‐latitude regions.

OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 3

1) Systems and cycles Some Definitions Transport and transformation processes within definite reservoirs: Carbon, Rock, Water Cycles Reservoir: (box, compartment: M in mass units or moles) An amount of material defined by certain physical, chemical, or biological characteristics that can be considered homogeneous: O2 in the atmosphere; carbon in living organic matter in the Ocean; ocean water in surface water masses. Flux: (F) The amount of material transferred from one reservoir to another per unit time (per unit area): The rate of evaporation of water from the surface ocean; the rate of deposition of inorganic carbon (carbonates in marine sediments); the rate of contaminant input to a lake or a bay Source: (Q) A flux of material into a reservoir Sink: (S) A flux of material out of a reservoir Budget: A balance sheet of all sources and sinks of a reservoir. If sources and sinks balance each other and do not change with time, the reservoir is in steady‐state (M does not change with time). If steady‐state prevails, then a flux that is unknown can be estimated by its difference from the other fluxes Turnover time: The ratio of the content (M) of the reservoir to the sum of its sinks (S) or sources (Q). The time it will take to empty the reservoir if there aren’t any sources. It is also a measure of the average time an atom/molecule spends in the reservoir. Cycle: A system consisting of two or more connected reservoir, where a large part of the material is transferred through the system in a cyclic fashion Feedback: All closed and open systems respond to inputs and have outputs. A feedback is a specific output that serves as an input to the system. Negative Feedback (stabilizing): The system’s response is in the opposite direction as that of the output. An example given in class is the increased reflection of solar radiation (albedo) from upper level clouds. Increased heat  evaporation  clouds  increased albedo  lowered incoming radiation  decreased overall heat. Positive Feedback (destabilizing): The system’s response is in the same direction as that of the output. An example given in class is the increased trapping of infrared radiation from lower level clouds. Increased heat  evaporation  clouds  increased I.R. trapping  increased overall heat. We also sent some time on the concept of residence time (a concept we will be using also in the second section of this course to explain the salt composition of seawater and biogeochemical cycles). Residence Time is a high probability that a certain fraction of a substance (atoms or molecules) forming the reservoir (M) will be of a certain age (mean age of the element when it leaves the reservoir). The residence time of water in the atmosphere is very short (~10‐20 days). The residence time of water in the Oceans is much longer (~4000 years). However, the residence time in different components of the atmosphere and oceans, and therefore the time of exchange between these different reservoirs, vary widely. OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 4

Figure 2: Time exchange for exchange of air and water between the atmosphere and ocean.

Advantages of Cycle Approach • Provides overview of fluxes, reservoir contents, and turnover time • Gives a basis for quantitative modeling • Helps to estimate the relative magnitudes of natural and anthropogenic fluxes • Stimulates questions such as: Where is the material coming from?, where is it going next? • Helps identify gaps in knowledge Disadvantages of Cycle Approach • Analysis, by necessity, superficial. Little or no insight into what goes inside the reservoir (“black box”) • Gives false impression of certainty. Often, at least one of the fluxes is derived from balance considerations (may be erroneous!) • Analysis based on average quantities that cannot always be easily measured because of spatial and temporal variations, as well as other factors.

2) Physical properties of water Water molecule: Dipole  Uneven charge  Hydrogen bonds! (DNA anyone?)  Higher energy requirement for change of state (solid to liquid, liquid to gaseous) than similar molecules. Make sure you can explain the figure below:

Figure 3: Melting and boiling temperatures for water and a series of molecules with similar chemical composition. OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 5

The structure of the water molecule thus leads to very high energy requirements for changes of state (Latent Heat), in particular for changes between liquid to gaseous state. In the figure below, the heat required to for changes in phase (state) are illustrated as horizontal lines. This demonstrates that heat has to constantly be supplied to water for his change of phase without, however, resulting in any change of temperature. Latent heat is just that, and change in heat without a change in temperature. Note the much more important heat requirement for vaporization (580 cal/gram) than for fusion (80 cal/gram). Also note that this heat transfer is reversible, meaning that 540 cal of heat is released to the atmosphere when 1 gram of water vapor condensed (rain) and 620 cal of heat is released when 1 gram of vapor solidifies (snow). Latent heat is thus an important component of the earth heat redistribution process (e.g. evaporation in inter‐tropical zones and condensation in mid‐ to high latitudes).

Figure 4: Heat and temperature changes in water across its phase change continuum. We also spent some time on the concept of Heat Capacity. Heat capacity is defined as the quantity of heat required to raise the temperature of 1 gram of a substance by 1°C. • More energy is required to raise the temperature of a substance with high heat capacity • At constant energy inputs, the substance with lower heat capacity will show a higher increase in temperature • High heat capacity substances can store large amount of energy.

We used the concept of heat capacity to explain major differences in temperature observed between continents and oceans. This was then applied to explain the three following figures.

Figure 5: Seasonal temperature curves at San Francisco (green) and Norfolk (blue). Both cities are located on the same latitude. Hence, differences are not due to solar radiation difference. OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 6

Instead, the moisture in the air in San Francisco transported from the Pacific through westerlies maintains the air temperature more stable over winter to summer seasonal changes. In contrast, the lack of moisture in the air at Norfolk (winds blow about land before reaching Virginia) is responsible for much larges seasonal changes in temperature.

Figure 6: Daily wind patterns in coastal regions. The winds are generated by pressure differences in the atmosphere, which are themselves the result of heat capacity differences between land and water.

Figure 7: Seasonal wind patterns in some coastal regions. The winds are generated by pressure differences in the atmosphere, which are themselves the result of heat capacity differences between land and water. Change in Temperature but not in Heat Adiabatic change: Water is slightly compressible (in the deep ocean where the weight of the water column induces a high pressure). This induces friction and thus higher kinetic energy  increase in temperature. In situ temperature: temperature measure on site (in situ). The utilization of in situ temperature can give the wrong impression that the water column in unstable (lighter water – warmer – in deep water masses). Temperature in deep water masses OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 7 needs to be corrected for the effect of pressure. The resulting temperature is called Potential Temperature. Change of Phase ­ Density During the transition from liquid to solid state, at the freezing point, the bond angle between oxygen and hydrogen atoms expands from about 105° to about 109°. This change allows ice to form a hexagonal crystal lattice. The space taken by 24 molecules in solid state could be occupied by 27 in the liquid state  Water expands about 9%! Ice has a density of 0.917 vs. ~1.000 g/cm3. Fresh water maximum density at ~4°C Density Density of seawater is affected by a combination of parameters: temperature (density when temperature), salinity (density when salinity), pressure (density when pressure). We can, however, remove the effect of pressure by using potential temperature (which itself is corrected for pressure). Hence, seawater density can be calculated as a function of both temperature and salinity: It is the density of a parcel of water of specific T and S that is brought up to the surface (no pressure effect!)

σT = ((1.02594/1.0000) – 1) x 1000 = 25.94 (No units!)

Figure 8: Sigma‐T values for water of different salinity (S) and temperature (T). To obtain the density of seawater you need to use σT and integrate it in the following equation: Density = [(σT × 10‐3) + 1] g/cm3. Note also that you can obtain seawater of similar densities by varying S and T. The examples circled in red, blue, and green show seawater of similar densities (within each color code) despite changes in parameters (e.g. as temperature increases, salinity has to increase for the water to maintain the same density). OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 8

Temperature­Salinity (TS) Diagrams In TS Diagrams (Figure 9), salinity (S) is represented on the x‐axis while temperature (T) is represented on the y‐axis. The sigma‐T (σT) lines indicate conditions of similar densities. The movement from the upper left to the lower right is a direction of increased density (the largest shift in density occurs when the movement is perpendicular to the σT lines).

Figure 9: TS Diagram. Note: A drop of 5°C in warm water (25°C) generates a greater increase in density than a similar cooling in cold water (5°C). In short, the more perpendicular to the σT lines the change is, the more intense the change in density.

Plotting actual values of T and S on such a TS Diagram will give you the number of water masses in the water column and their specific T and S characteristics.

Figure 10: Depth profiles of temperature and salinity for a station in the North Atlantic. OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 9

Figure 11: TS Diagram for the station shown in Figure 10. Surface water is the lightest and thus appears at the top of the diagram. Heaviest water appears at the bottom of the diagram. As stated in the figure, every “bend” in the curve denotes a water mass that does not mix with water above and below.

Caballing: When two water masses of similar densities merge (points a and b in Figure 12 below), the combination of their temperatures and salinities results in densification  vertical advection of water.

Figure 12: TS Diagram showing the effect of cabaling. The water in c (a 50:50 mixture of water a and b) is denser than the two original water masses.

OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 10

Light Autotrophs – with a few exceptions – depend on energy from sunlight. As do land plants, marine plants use chlorophyll and other pigment to capture the visible light from the sun to perform photosynthesis. As solar radiation strikes the surface of the Ocean, a large fraction of it is reflected back to the atmosphere (dependent on the angle of the sun’s rays and the smoothness of the water surface). The amount that enters is ultimately absorbed by water molecules (~65% of visible light is absorbed within 1 m depth!):  Absorbed energy manifests itself as heat (elevating the temperature of the surface water)

Figure 13: Spectrum of light absorption with respect to water depth. Absorption is greatest at longer wavelength. In clear water, only ~1% of surface energy remains at 100 m (in coastal waters with lots of particles, light doesn’t penetrate more than a few meters).

3) Heat budget of the earth Based on the temperature of the sun and the earth, and on physical laws of radiation (Stefan­ Boltzmann’s and Wien’s laws), the sun emits radiation mostly in the visible to ultra violet (UV), whereas the earth emits mostly in the infrared (IR) spectrum. These differences in radiation wavelengths are crucial to explain the earth actual vs. theoretical average temperature. Based on the solar radiation that earth receives per unit area, the theoretical temperature of the earth should be ‐18°C. But the actual temperature of earth is much higher than that (+15°C). The +33° difference can be explained by what is called the greenhouse effect. In short, gases in the atmosphere (e.g. water, CO2, CH4, CFCs, N2O) are relatively transparent to short wave radiation (meaning they let these pass through), whereas they absorb longer wavelength radiation. Figure 14 illustrates this process. The substantial absorption of IR back radiation by the earth atmosphere thus permits the earth to retain heat within in fluid envelope and thus warm beyond the theoretical value determined by solar radiation alone (Figure 15). OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 11

Figure 14: Proportion of light absorption in the earth atmosphere with respect to radiation wavelength. There are two “window” of transparency. The atmosphere is nearly transparent to visible and near UV radiation (majority of solar radiation), whereas it absorbs strongly in the UV (from O3) and near IR (water, CO2, CH4, CFCs, N2O). The second window is in the IR and permits some long wavelength radiation to escape the earth’s atmosphere.

Figure 15: Warming of the earth by solar radiation alone (left) and through the combination of solar radiation and greenhouse effect (right).

Figure 16. Although average temperatures vary seasonally and spatially, the earth’s overall T changes only slightly over the years  “must” return to space the same amount of energy it absorbed. Total energy input = 100 units (per unit time). Note: Albedo is the fraction of the sun radiation that is reflected back to space without being incorporated in the heat budget of earth. OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 12

Two conditions/processes affect the amount of solar radiation (and therefore the incoming energy source) different regions of the earth receive: Sphericity and Seasonality. Sphericity (below): If the Earth were a disk with its surface perpendicular to the rays of sunlight, each point on it would receive the same amount of radiation. However, the Earth is a sphere and its surface tilts with respect to the incoming rays of energy with the regions furthest away aligned in parallel to the radiation and thus receiving no energy at all.

Seasonality (right): The Earth’s axis is tilted in relation to the plane of the ecliptic (the Sun‐Earth plane that cuts through both centers). Higher latitudes thus receive incoming radiation with different angles through the seasons (very low angle in the winter and close to 90° in the summer).

The effect of seasonality and sphericity is illustrated in the figure below where short wave incoming radiation is maximum in high latitude summers (july in the north hemisphere and January in the south hemisphere) and constant at the equator.

OCNG 251 (Dr. P. Louchouarn) – Study Guide #1 13

Because of these geographical and temporal variations in short wave incoming radiation, long wave radiation also show similar variations (IR). The figure below illustrates these geographical and temporal differences.

Figure 17. Long wave radiation for the earth. Left hand panel: January. Right hand panel: June.

These figures show graphically that the there is an equator to pole change in incoming radiation  ΔHeating as well as a seasonal change in incoming radiation  ΔHeating.

The earth heat budget is thus apparently unbalanced with the inter‐tropical region apparently continuously gaining heat and the high latitude regions apparently continuously losing heat (Figure 18). However, because the inter‐tropical zones do not heat continuously or the high latitude regions do not freeze continuously, then heat transfer from zones of heat surplus to the zones of heat deficit must occur. This is the founding condition that leads to circulation patterns in the fluid envelopes of earth (atmosphere and ocean).

Figure 18 Short wave (Qs) minus long wave (Qb) radiation for the earth.