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Chapter 5. The Hydrosphere as Microbial Habitat

Chapter 5. The Hydrosphere as Microbial Habitat

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58



Geomicrobiology





90°



180°



90°



Arctic Ocean

70°



70°



45°



45°

Atlantic

Ocean



Pacific

Ocean









Indian

Ocean



45°



45°



FIGURE 5.1 Oceans of the world.

Shore

Shelf



tine



Con

ntal



Ocean basin



e



slop



Sea mount



Ris



e



Continental margin



Abyssal plain

Sea floor



Schematic representation of a profile of an ocean basin.



dropped, and when settling it abrades the slope. Marine canyons may also be cut by slumping of an

unstable sediment deposit on a portion of the continental slope and the consequent abrasion of the

slope. Occasionally, the continental slope may be interrupted by a terraced region, as in the case of

Blake Plateau off the southern Atlantic coast of the United States. This particular shelf is ∼302 km

wide and drops gradually from a depth of 732 to 1100 m over this distance. It was gouged out of the

continental slope by the northward flowing Gulf Stream.

At the foot of the continental slope lies the continental rise, consisting of accumulations of sediment carried downslope by turbidity currents. Such deposits may extend for 100 km or more from

the foot of the continental slope. The continental rise may form fanlike structures in some places

and wedges in others. An idealized profile of a continental margin is shown in Figure 5.2.

The ocean basin takes up 80% of the ocean area. Its floor, far from being a flat expanse, as once

believed by some, often exhibits a rugged topography. Submarine mountain ranges cut by fracture

zones and rift valleys stretch over thousands of kilometers as the midocean ridge systems where the

new ocean floor is created (see Chapter 2). Elsewhere, somewhat more isolated submarine mountains, some of which are active and others dormant volcanoes, dot the ocean floor. Some of the

seamounts have flattened tops and have been given the special name of guyots. Some of the flattened

tops of seamounts, especially in the Pacific Ocean, reach surface waters at depths of 50–100 m



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59



where the temperature is ∼21°C. In these positions, the flat tops may serve as substratum for colonization by corals (coelenterates) and coralline algae, which then form atolls and reefs.

Covering the ocean floor almost everywhere are sediments. They range in thickness from 0

to 4 km, with an average thickness of 300 m. Their rate of accumulation varies, being slowest in

midocean (<1 cm per 103 years) and fastest on continental shelves (10 cm per 103 years). These

rates may be even greater in some inland seas and gulfs (e.g., 1 cm per 10–15 years in the Gulf of

California and 1 cm in 50 years in the Black Sea). In some regions of the deep ocean, the sediments

consist mainly of deposits of siliceous and calcareous remains of marine organisms. The siliceous

remains are derived from the frustules of diatoms (algae) and the support skeleton and spines of

radiolarians (protozoa). The calcareous remains are derived from the tests of foraminifera (protozoa), carbonate platelets from the walls of coccolithophores (algae), and shells from pteropods (mollusks). Diatomaceous oozes predominate in colder waters (e.g., in the North Pacific between 40°

and 70° north latitude and 140° west to 145° east longitude; Horn et al., 1972). Radiolarian oozes

predominate in warmer waters (e.g., in the North Pacific between 5° and 20° north latitude and 90°

and 180° west longitude; Horn et al., 1972). Calcareous oozes are found mainly in warmer waters

on ocean bottoms no deeper than 4550–5000 m (e.g., in the North Pacific between 0° and 10° north

latitude and 80° and 180° west longitude; Horn et al., 1972). At greater depths, the CO2 concentration in the water is high enough to cause dissolution of carbonate unless the structures are enclosed

in a protective membrane.

Other vast areas of the ocean floor are covered by clays (red clay or brown mud), which are

probably of terrigenous origin and washed into the sea by rivers and general runoff from continents

and islands and carried into the ocean basins by ocean currents, mudflows, and turbidity currents.

At high latitudes in both hemispheres, particularly on and near continental shelves, ice-rafted sediments are found. They were dropped into the ocean by melting icebergs, which had previously separated from glacier fronts that had picked up terrigenous debris during glacial progression. Except

for ice-rafted detritus, only the fine portion of terrigenous debris (clays and silts) is carried out to

sea. The clay particles are defined as having a diameter <0.004 mm, and the silt particles as having

a size range of 0.004–0.1 mm in diameter. Figure 5.3 shows the appearance of some Pacific Ocean

sediments under the microscope.



5.1.2 OCEAN IN MOTION

A significant portion of the ocean is in motion at all times (Williams, 1962; Bowden, 1975). The

causes of this motion are (1) wind stress on the surface waters, (2) Coriolis force arising from

the rotation of the Earth, (3) density variation of seawater resulting from temperature and salinity

changes, and (4) tidal movement due to gravitational influences on the water exerted by the sun and

the moon. Surface currents (Figure 5.4) are prominent in regions of prevailing winds, such as the

trade winds, which blow from east to west at ∼20° north and south latitudes; the westerlies, which

blow from west to east between 40° and 60° north and south latitude; and the easterly polar winds,

which blow in a westerly direction south of the Arctic Ocean. The effect of these winds, together

with the deflecting influence of the continents and the Coriolis force, is to set up surface circulations

in the form of gyrals between the north and south poles in each major ocean. They are the North

Subtropical Gyral (large), the North Tropical Gyral (small), the South Tropical Gyral (small), the

South Subtropical Gyral (large), and the Antarctic Current that circulates around Antarctica from

west to east (Figure 5.4A). Thus the Gulf Stream, together with the Canary Current and the North

Equatorial Current, is part of the North Subtropical Gyral of the Atlantic Ocean (Figure 5.4B). The

flow rates of the waters in these gyrals and segments of them are different. The flow rate of the water

in the Gulf Stream is fastest of any surface current—250 cm s−1. Other currents have flow rates that

are mostly in the range of 25–65 cm s−1.

Meanders in the Gulf Stream in the Atlantic and the Kuroshio Current in the Pacific Ocean

may give rise to the so-called rings—small closed current systems that may measure as much as



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FIGURE 5.3 Microscopic appearance of marine sediments (×1750). (A) Atlantic sediment showing coccoliths (CaCO3) (arrows) and clay particles. (B) Atlantic sediment showing diatom frustules (SiO2) and other

debris. (C) Pacific sediment showing a centric diatom frustule (SiO2) and other debris. (D) Pacific sediment

showing fragments of radiolarian tests (SiO2).



300 km in diameter and may have a depth as great as 2 km. Such rings may move 5–10 km per

day. The chemical, physical, and biological characteristics of the water enclosed in a ring may be

significantly different from those of the surrounding water. A slow exchange of solutes and biota as

well as heat transfer may take place across the boundary of a ring. Rings thus constitute the means

of nutrient transport from ocean currents. The rings may ultimately rejoin the current that spawned

them (Gross, 1982; Richardson, 1993; Ring Group, 1981). More recently, anticyclones have been

reported to arise from the Gulf Stream in addition to the rings, and meddies from the north of the

Strait of Gibraltar (Richardson, 1993). Although rings have a cold water core surrounded by a warm

water layer and counterclockwise rotation, anticyclones have a warm water core surrounded by

colder water and clockwise rotation. Meddies have a core that is more saline than the surrounding

ocean water and a clockwise rotation. Collectively, these formations are known as ocean eddies.

Deep water is also in motion. Its movement appears to be caused by slow diffusion resulting from

density differences of water masses through broad zones in the oceans. Some of the deep currents

that have been measured in the Atlantic Ocean have a velocity between 1 and 2 cm s−1 (Dietrich and

Kalle, 1965, pp. 399, 407; Gross, 1982). Occasional short bursts in velocity may occur. The bottom

current movement is influenced by bottom topography.

Deep water may rise toward the surface in a process called upwelling. This results from the

moving apart of two surface water masses, causing the deep water to rise to take the place of the

divergent waters. The moving apart of the water masses is called divergence (Williams, 1962).



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Northern Boundary

Polar

winds



North Subpolar Gyral



rly

ste

We inds

w



North

Subtropical

Gyral



de

Tra ds

win



North Tropical Gyral



Equator



Equator



South Tropical Gyral



South

Subtropical

Gyral



Westerly



Tra

win de

ds



Antarctic Current



winds

(A)

40°°



60°°



80°° 100°° 120°° 140°° 160°° 180°° 160°° 140°° 120°° 100°° 80°°



60°°



40°°



20°°



0°°



20°° 40°°



nd



20°°



80°°



ee



nla



80°°



Gr



No



E



70°°



rw

ay



70°°



Subarctic

io



North Equatorial



Countercurrent



torial



0°°



Guinea



az

il

Br



Peru



E Australia



20°°



kla



nd



40°°



Fa

l



W Australia



lha

Agu



Equa



South Equatorial



la

ue



40°°



20°°



ng



s



South Equatorial



uth



Be



20°°



40°°



So



Countercurrent



North Equatorial



0°°



ia



rn



20°°



tem

Stream Sys

f

Glu

North Equatorial



lifo



h

ros

Ku



Ca



North Pacific



Can



40°°



ary



ka



O



or



as



Al



s

ya



ad



br



o

hi



60°°



Irminger



La



60°°



West Wind Drift



60°°



West Wind Drift



West Wind Drift



60°°

70°°



70°°

20°°



40°°



60°°



80°° 100°° 120°° 140°° 160°° 180°° 160°° 140°° 120°° 100°° 80°°



60°°



40°°



20°°



0°°



20°° 40°°



(B)



FIGURE 5.4 Oceanic surface currents. (A) Schematic representation of the prevailing winds and their effects

on the surface currents of an imaginary rectangular ocean. (B) Average surface currents of the world’s oceans.

(From Williams J, Oceanography, Little, Brown, Boston, MA, 1962. With permission. [Jerome Williams is

deceased.])



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TABLE 5.1

Some Constituents of Seawater (àg L1)

Major Constituents

Cl

Na

Mg

S (SO4)

Ca

K

Br

C (CO3, HCO3)

B



Minor Constituents

1.9 ì 107

1.1 × 107

1.3 × 106

9.0 × 105

4.1 × 105

3.9 × 105

6.7 × 104

2.8 × 104

4.5 × 103



Si

N

Li

Rb

P

I

Ba

Mo

Zn

Ni



3 × 103

6.7 × 102

1.7 × 102

1.2 × 102

90

60

20

10

10

7



Cu

Fe

U

As

Mn

Al

Co

Se

Pb

Ra



3

3

3

2.6

2

1

0.4

9 × 10−2

3 × 10−2

1 × 10−7



Source: Marine Chemistry. A Report to the Marine Chemistry Panel of the Committee

of Oceanography, National Academy of Sciences, Washington, DC, 1971.



Upwelling of deep water may also result when winds blow large surface water masses away from

coastal regions (Smith, 1968). Deep water is relatively rich in mineral nutrients, including nitrate

and phosphate, and thus, upwelling is of great ecological significance because it replenishes biologically depleted nutrients in the surface waters. Regions of upwelling are therefore very fertile. An

important region of upwelling in the eastern Pacific Ocean occurs off the coast of Peru. A disturbance in the surface water circulation in the southern Pacific can result in failure of upwelling in this

region (El Niño) and can spell temporary disaster for the fisheries of the area.

When a dense surface water mass meets a lighter water mass, a convergence occurs, and the

heavier water will sink to a level where it meets water of the same density. This phenomenon is

important because the denser, sinking surface water will carry oxygen to the deep waters. Important

convergences in the oceans occur at high latitude in both the hemispheres.



5.1.3



CHEMICAL AND PHYSICAL PROPERTIES OF SEAWATER



Seawater is saline. Some important chemical components of seawater, listed in decreasing order of

concentration, are presented in Table 5.1 (Marine Chemistry, 1971). Of these components, chloride

(55.2%), sodium (30.4%), sulfate (7.7%), magnesium (3.7%), calcium (1.16%), potassium (1.1%), bromide (0.1%), strontium (0.04%), and borate (0.07%) account for 99.5% of the total salts in solution

(percent, in weight per volume). Because these components generally occur in constant proportions

relative to one another in true ocean waters, it has been possible to estimate salt concentrations in

seawater samples by merely measuring chloride concentration. The chloride concentration in grams

per kilogram (chlorinity, Cl) is related to the total salt concentration (salinity, S) in grams per kilogram by the relationship

S(‰) = 0.030 + 1.8050Cl (‰)*



(5.1)



The salinity so determined is an estimate of the total amount of solid material in a unit mass of

seawater in which all carbonate salts have been converted to oxides and all bromide and iodide

have been replaced by chloride, and in which all organic matter has been completely oxidized.

* The symbol “‰” represents parts per thousand or grams per kilogram. Equation 5.1 was amended to S(‰) = 0.030 +

1.80655Cl (‰) by UNESCO in 1969.



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For reference purposes, the salinity of standard seawater has been taken as 34.3‰. The actual

salinity of different parts of the world oceans can vary from <33.4‰ to almost 36‰ (Dietrich and

Kalle, 1965, p. 156).

Accurate titrimetric chlorinity determinations for determining seawater salinity are cumbersome and require skill. Nowadays, determinations of salinity are made using a salinometer, which

measures electrical conductivity of seawater. Such measurements are translated into practical salinity (S) using the following relationship formulated and adopted by UNESCO/CES/SCOR/IAPSO

Joint Panel on Oceanographic Tables and Standards in 1978:

1/ 2

3/2

2

5/2

S ϭ 0.008 Ϫ 0.1692 K15

ϩ 25.3851K15 ϩ 14.0941K15

Ϫ 7.0261K15

ϩ 2.7081K15



(5.2)



where K15 represents the ratio of electrical conductivity of a seawater sample at 15°C and 1 standard

atmosphere of pressure to that of a standard KCl solution containing 35.4356 g of KCl in 1 kg of

solution (see http://www.salinometry.com/content/view/18/31/). The value of K15 for seawater having a salinity of 35‰ is exactly unity. Practical salinity as presently defined is dimensionless.

Table 5.2 lists the salinities, determined from chlorinity measurements, of some different marine

waters as well as those of some saline lakes. It must be pointed out that whereas the Great Salt

Lake in Utah has a salt composition that is qualitatively similar to that of oceans, some other inland

hypersaline water bodies, such as the Dead Sea at the mouth of the Jordan River, have a different

salt composition. In the Dead Sea, the dominant cations and their respective concentrations are in

descending order: Mg (∼44 g L−1), Na (40 g L−1), Ca (17 g L−1), and K (7.5 g L−1), and the dominant anions and their respective concentrations are Cl (225 g L−1) and Br (5.5 g L−1) (Nissenbaum,

1979).

Although some portions of the salts in seawater derive from the runoff from the continents and

the weathering of minerals in the surficial sediments, a most important contribution to seawater

solutes is made by hydrothermal discharges from vents at the midocean spreading centers. These

discharges are the consequence of seawater penetration deep into the porous basalt (down to depths

of 1–3 km) beneath the ocean floor, where the seawater then reacts with constituents of the basalt

in various ways. The reactions include the reduction of seawater sulfate to hydrogen sulfide by ferrous iron in the basalt. They also include the removal of magnesium from seawater as magnesium

hydroxide and the incorporation of seawater calcium into minerals such as plagioclase to form new

aluminosilicate minerals such as clinozoisite or Ca-rich amphibole, accompanied by the generation of acidity. In the case of calcium incorporation into plagioclase, this can be illustrated by the

reaction

3CaAl2Si2O8 + Ca2+ + 2H2O → 2Ca2Al3Si3O12(OH) + 2H+



(5.3)



TABLE 5.2

Salinities (‰) of Some Marine Waters and Salt Lakes



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Water Body



Salinity(‰)



References



Gulf of Bothnia

Baltic Sea

Black Sea

Mediterranean Sea

Red Sea

Dead Sea

Great Salt Lake

Ocean bottoms



2–6

6–17

16–18

37–39

40–41

320

320

34.66–34.92



Smith (1974)

Smith (1974)

Smith (1974)

Bowden (1975)

Bowden (1975)

ZoBell (1946)

Zobell (1946)

Defant (1961)



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The resultant acidity is the cause of leaching of some other components from the basalt, such as

hydrogen sulfide from pyrrhotite and base metals from some other basalt minerals. All these reactions are possible because of high hydrostatic pressure (HP) exerted on the solution in the basalt at

these depths and because of the high temperature (∼350°C) from heat diffusion from the underlying

magma chambers into the reaction zone of the basalt. The resultant chemically altered seawater is

forced upward by HP through porous channels and fissures in the basalt. It is ultimately discharged

as a hydrothermal solution from vents at the spreading centers into the overlying seawater and

mixed with it (Bischoff and Rosenbauer, 1983; Edmond et al., 1982; Seyfried and Janecky, 1985;

Shanks et al., 1981) (see also Chapters 2, 17, and 20). These reactions contribute significantly to the

stability of seawater composition.

Seawater contains a pH-buffering system that consists of bicarbonate and carbonate ions, borates,

and silicates. The carbonate plus bicarbonate ions constitute 0.35% of the solutes in seawater.

Together, these buffers keep the pH of seawater in the range of 7.5–8.5. Surface seawater pH tends

to fall into a narrow range of 8.0–8.5. At depth, seawater pH may approach 7.5. To a major extent,

the variation in pH of seawater with depth may be related to oxygen utilization during respiration

by marine organisms, which results in CO2 production from organic carbon. To a lesser extent, it

may be related to carbonate minerals (e.g., CaCO3) dissolution (Park, 1968). Figure 5.5 illustrates

the changes in pH with depth at one particular station in the Pacific Ocean.

Because of the alkaline pH and elevated Eh of seawater, nutritionally available iron appears to be

a limiting micronutrient for bacterioplankton and phytoplankton (Tortell et al., 1999; Hutchins et al.,

1999; Gelder, 1999). This is because iron under these conditions will be ferric and, unless complexed, will predominate in the form of insoluble hydroxide, oxyhydroxides, and oxides. Various

bacteria and algae release ligands (siderophores) that complex with ferric iron and thus keep it

in solution at the alkaline pH of seawater and make it nutritionally available (Tortell et al., 1999;

Hutchins et al., 1999; Martinez et al., 2000). Growth by some siderophore-producing marine bacteria in iron-limited waters can be stimulated by exogenous siderophores (Guan et al., 2001). In at

least some parts of the oceans, it is possible to stimulate growth of phytoplankton and bacterioplankton by fertilizing the ocean water with iron (Coale et al., 1996; Church et al., 2000; Arrieta

et al., 2004). Growth stimulation of phytoplankton in the ocean by iron fertilization might offer

a means of significant enhancement of CO2 sequestration in the ocean and, thus, have a positive

pH (25°C)

7.4 7.6



7.8 8.0 8.2



0



Depth (m)



2000



4000



6000

−0.6 −0.4 −0.2 −0.0 +0.2

∆pH



Vertical profile of pH at station 54°46′N, 138°36′W in the northeastern Pacific Ocean. (Adapted

from Park PK, Science, 162, 357–368, © 1968 by the American Association for the Advancement of Science.

With permission.)



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effect on climate control. It might even enhance the bioproduction of silica by growth stimulation

of marine diatoms, an important food source for some plankton feeders, for instance, in the seas

around Antarctica (Southern Ocean). However, theoretical considerations and results from field

experiments examining these possibilities suggest that the extent of sequestration of carbon and its

transfer to the ocean floor as particulate organic carbon (POC) is not great enough to have a significant impact on climate control (Liss et al., 2005; Zeebe and Archer, 2005; Brzezinski et al., 2005;

Arrieta et al., 2004; Buessler et al., 2004; Buessler and Boyd, 2003; Bakker, 2002).

The salts dissolved in seawater impart a special osmotic property to it. The osmotic pressure

of seawater is of the order of magnitude of the internal pressure of bacterial cells or the cell sap of

eukaryotic cells. At a salinity of 35‰ and a temperature of 0°C, seawater has an osmotic pressure of

23.37 bar (23.07 atm), whereas at the same salinity but at 20°C it has an osmotic pressure of 25.01 bar

(24.96 atm). Clearly then, the osmotic pressure of seawater is not deleterious to living cells.

With increasing depth in the water column, HP becomes a significant factor in the life of microbes

and other forms of life in the sea. On average, the HP in the open ocean increases ∼1 atm (1.013 bar)

for every 10 m of depth. Related to the weight of overlying water at a given depth, HP in the oceans

ranges from 0 to more than 1000 atm (1013 bar). Thus, the highest pressures occur in the deep ocean

trenches. Among the marine fauna, some members are adapted to live only in surface waters, others

at intermediate depths, and still others at abyssal depths. Generally, none are known to live over

the entire depth range of the open ocean. Although microorganisms such as bacteria appear to be

more adaptable to changes in HP, facultative (pressure tolerant) and obligately barophilic (pressurerequiring) bacteria are known (see also Section 5.1.5).

Salinity and temperature affect the density of seawater. At 0°C, seawater having a salinity of

30–37‰ has a corresponding density range of 1.024–1.030 g cm−3. A variation in seawater density

due to variation in salinity is one cause of water movement in the ocean, because denser water will

sink below lighter water (convergence), or conversely, lighter water will rise above denser water

(upwelling). The following processes may cause changes in salinity and, therefore, density: (1) dilution of seawater by runoff or less-saline water; (2) dilution by rain or snow; (3) concentration through

surface evaporation; (4) freezing, which excludes salts from ice and thus leaves any residual, unfrozen water more saline; or (5) thawing of ice, which dilutes the already existent saline water.

As already stated, variation in salinity of seawater is not the sole cause of variation in density.

The other important cause of density variation of seawater is temperature. Unlike freshwater, whose

density is greatest at 4°C (Figure 5.6B), seawater with a salinity of 24.7‰ or greater has maximum

density at its freezing point, that is, 0°C (Figure 5.6A). A body of freshwater thus freezes from its

surface downward because freshwater at its freezing point is lighter than at a temperature of 4°C.

Ocean water in the Arctic or Antarctic seas also freezes from the surface downward, but in this

instance because ice, which excludes salts as it forms from seawater, is lighter than the seawater and

will thus float on it.

The temperature of seawater ranges from about −2°C (the freezing point at 36‰ salinity)

to +30°C, in contrast to the temperature of air over the ocean, which can range from −65 to +65°C.

The narrower temperature range for seawater can be related to (1) its heat capacity, (2) its latent

heat of evaporation, and (3) heat transfer from lower to higher latitudes by surface currents in both

hemispheres. The major source of heat in the ocean is solar radiation. More than half the surface

waters of the ocean are at 15–30°C. Only 27% of the surface waters are below 10°C. From ∼50°

north latitude to 50° south latitude, the ocean is thermally stratified. In this range of latitudes, the

seawater temperature below the depth of ∼1000 m is below 4°C (deep water). At depths from ∼300

to 1000 m, the temperature drops rapidly with increasing depth. The zone of this rapid temperature

change is called the thermocline. Its thickness and position vary with geographic location and season of the year. Above the thermocline lies the warm surface water, the mixed layer,

r which is extensively agitated by wind and water currents and thus exhibits relatively little temperature change with

increasing depth.



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Salinity (% )



40



24.7%

20



+4



+2



0

Temperature (°C)



−2



−4



2



4

Temperature (°C)



6



8



(A)

1.00000

98



Density



96

94

92

90

88

0.99986



0



(B)



FIGURE 5.6 Density relationships in seawater and freshwater. (A) Relationship of seawater salinity to freezing point. Symbols: open circles, temperature of maximum density at a given salinity; closed circles, freezingpoint temperature at a given salinity. Note that above a salinity of 24.7‰ seawater freezes at its maximum

density as its temperature at maximum density cannot be lower than its freezing point. (B) Relationship of

freshwater density (in g cc−1) to temperature. Data points from chemically pure water are shown. Note that in

the case of freshwater, its density at its freezing point is lower than its density at 4°C.



At latitudes higher than 50°N and 50°S, seawater is not thermally stratified. The waters around

Antarctica, being cold (−1.9°C) and hypersaline (34.82‰) owing to ice formation, are hyperdense

and thus sink below warmer, less-dense water to the north and flow northward along the bottom of

the ocean basin. This is an example of convergence. Similarly, Atlantic waters from the subarctic

region, having a temperature in the range of 2.8–3.3°C and a salinity in the range of 34.9–34.96‰,

sink and flow southward at near-bottom or bottom levels of the ocean. Because the Arctic Ocean

bottom is separated from the other oceans by barriers such as the shallow Bering Strait in the case

of the Pacific Ocean and a shallow ridge in the case of the Atlantic Ocean, it does not influence the

water masses of the Pacific and Atlantic Oceans directly. Other convergences occur in the world’s

oceans in both hemispheres because of the interaction of waters of different densities. In these

instances, the heavier waters sink to lesser depths because they have lower densities than the heavier

waters at high latitudes.

The water convergences alluded to above help to explain why generally ocean water is oxygenated at all depths (Figure 5.7; Kester, 1975). Of all ocean waters, only some coastal and near-coastal



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Oxygen (µ mol kg−1)



0



100



200



Depth (km)



0



1



1



2



300



NE Atlantic

29.5°N 56.5°W

Kester (1972)



2



100



0



200



Indian Ocean

13°S 75°E

Rakestraw

(1964)



1



3



3



4



4



4



Bottom



200



300



NE Pacific

38.0°N 124.8°W

Culberson and

Pytkowicz (1970)



2



3



5



100



Bottom



5



Vertical distribution of oxygen in the ocean. Profiles from three ocean basins. (From Kester

DR, Chemical Oceanography, Academic Press, New York, 1975. With permission. Copyright by Elsevier.)



waters (e.g., estuarine waters; Cariaco Trench) may, as a result of intense biological activity, be

devoid of oxygen at depth. At some sites, intense biological activity is sometimes the direct result

of pollution caused by human beings. Surface waters of the open ocean tend to be saturated with

oxygen because of oxygenation by the atmosphere and, equally important, by the photosynthetic

activity of the phytoplankton. Oxygenation by phytoplankton can occur to depths of ∼100 m (200 m

in exceptional cases), where light penetration is 1% of the surface illumination. Seawater salinity of

34.352‰ is saturated at 5.86 mL or 8.40 mg of oxygen per liter at 760 mm Hg and 15°C. The higher

the salinity and the higher the temperature, the lower the solubility of oxygen in seawater.

Starting at the top of the water column, the oxygen concentration in seawater will at first decrease

with depth, owing mainly to oxygen consumption by the respiration of living organisms (Figure 5.7). Because many life forms in the oceans tend to be concentrated in the upper waters, oxygen

concentration will fall to a minimum at ∼600–900 m of depth, where respiration (oxygen consumption) by zooplankton and other animal forms as well as bacterioplankton occurs but not photosynthesis (oxygen production) by phytoplankton. Below this depth, because of rapidly decreasing

biological activity, the oxygen concentration may at first increase once more and then slowly

decrease again toward the bottom. Bottom water, however, may still be half-saturated with oxygen

relative to surface water. This oxygen is not supplied by in situ photosynthesis, which cannot occur

in the absence of light at these depths, nor is it the result of significant oxygen diffusion from the

atmosphere to these depths. As previously indicated, the oxygenated waters at depth derive from the

antarctic and subarctic convergences. The oxygen-carrying waters from the antarctic convergences

flow northward along the bottom and at intermediate depths of the ocean basins, whereas the waters

from subarctic convergence in the Atlantic flow southward at more intermediate depths. The oxygen

content of these waters is only slowly depleted because of the low numbers of oxygen-consuming

organisms in these deep regions of the oceans and the low rate of oxygen consumption in the upper

sediments.

Photosynthetic activity of the phytoplankton is dependent on the penetration of sunlight into

the water column because phytoplankton derives its energy almost exclusively from sunlight. It

has been shown that light absorption by pure water in the visible range between 400 and 700 µm

increases greatly toward the red end of the spectrum. It has also been shown that 60% of the light

that penetrates transparent water is absorbed at a depth of 1 m. And 80 and 99% of the same light is

absorbed at depths of 10 and 140 m, respectively. In less transparent coastal water, 95% of the light



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68



Geomicrobiology



may have been absorbed at 10 m. Although the photosynthetic process of phytoplankters can use

light over the entire visible spectrum, action spectra show peaks in the red and blue ends of the spectrum, where chlorophylls absorb optimally. Accessory pigments, such as carotenoids, absorb light

at intermediate visible wavelengths. Clearly, light penetration limits the depth at which phytoplankton can grow. This depth is ∼80–100 m on average (200 m maximally) and often much less in less

transparent waters. Two exceptions have been noted, however. One was seen off the northern border

of San Salvador Island in the Bahamas, where crustose coralline algae ((Rhodophyta) were growing

attached to rock at a depth of 268 m, observed from a submersible. At this location, the light intensity was only ∼0.0005% of that at the surface (Littler et al., 1985). The other exception was noted in

the Black Sea. Here, the photosynthetic sulfur bacterium Chlorobium phaeobacteroides was found

to grow in a chemocline at a depth of 80 m, where light transmission from surface irradiance has

been calculated to be 0.0005% (Overmann et al., 1992), as at the station at San Salvador Island.

The water layer from the ocean surface to the depth below which photosynthesis cannot take place

constitutes the euphotic zone. Zooplankton and bacteria, except for cyanobacteria, may abound to

somewhat lower depths than phytoplankton (∼750 m), being scavengers and able to feed on dying

and dead phytoplankters and their remains in the process of settling.



5.1.4



MICROBIAL DISTRIBUTION IN WATER COLUMN AND SEDIMENTS



Microbial distribution in the open oceans is not uniform throughout the water column (Figure 5.8).

Factors affecting this distribution are energy, carbon, nitrogen, and phosphorus limitations (Wu

et al., 2000) and also temperature, HP, and salinity. Accessory growth factors, such as vitamins, may



0



Bacteria cm−3 × 100

1

2

3

4

5



0



mg L−1

Light 0 20 40 60 80 100%

0.1 0.2 0.3 0.4 0° 10° 12° 14° 16° 18° 20°C



Light



0

25

s



50



tom



Dia



NO3

Ba



ct



PO4



100



Temperature



Depth (m)



75



ia

er



150



200



9 × 103 − 108 bacteria per gram of mud

Sea bottom



FIGURE 5.8 Vertical distribution of bacteria (number per cubic centimeter of water), diatoms (number per

liter of water), PO4, NO3 (milligrams per liter), light, and temperature in the sea based on average results at

several different stations off the coast of Southern California. (Reproduced from ZoBell CE, Sci Mon., 55,

320–330, © 1942 by the American Association for the Advancement of Science. With permission.)



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