Hydrogeochemistry Fundamentals and Advances, Environmental Analysis of Groundwater

By Viatcheslav V. Tikhomirov

Water is the Earth's most precious resource.  Until recent years, water was often overlooked as being overly abundant or available, but much has changed all over the world.  As climate change, human encroachment on environmental areas, and deforestation become greater dangers, the study of groundwater has become more important than ever and is growing as one of the most important areas of science for the future of life on Earth.

This three-volume set is the most comprehensive and up-to-date treatment of hydrogeochemistry that is available. The first volume lays the foundation of the composition, chemistry, and testing of groundwater, while volume two covers practical applications such as mass transfer and transport. Volume three, which completes the set, is an advanced study of the environmental analysis of groundwater and its implications for the future.

This third volume focuses more deeply on the analysis of groundwater and the practical applications of these analyses, which are valuable to engineers and scientists in environmental science, groundwater remediation, petroleum engineering, geology, and hydrology. Whether as a textbook or a reference work, this volume is a must-have for any library on hydrogeochemistry.

• Language: English
• Publisher: Wiley
• Release date: Feb 27, 2018
• ISBN: 9781119160519

Book preview

Chapter 1

Hydrosphere

The amount of moisture on Earth and its isotopic composition support a suggestion that our water emerged in the process of the formation of the Earth as a planet at the earliest stages of the Solar System evolution. According to the accretion theory, the major source of moisture on Earth could have been cosmic material from two possible basic sources: asteroids and comets, 70–80% of whose volume is water. However, the first measurements showed that in carbonaceous chondrites the D/H isotope ratio is equal to (1.4 ± 0.1) × 10−4 and in six comets from the Oort’s cloud it is (2.96 ± 0.25) × 10−4 (Figure 1.1) (Pinti, 2005; Van Kranendonk, 2012). The closeness of the former value to the D/H ratio in the ocean (1.558 ± 0.001) × 10−4 suggests that 90% of moisture volume on Earth had come from the asteroid belt positioned between the orbits of Mars and Jupiter, and only 10%, from comets. In 2011, the entire game had been upset by the isotopic composition of moisture in a comet from the Jupiter family, 103P/Hartley 2 (Hartogh et al., 2011). The D/H moisture ratio value for this comet was found to be equal to (1.61 ± 0.24) × 10−4, i.e., close to the ocean value. This discovery substantially broadened the range of the outer space material capable of having served the source of moisture on Earth and now includes the belt of comets positioned directly beyond the planet Neptune (30–50 a.u.). These data satisfy the requirements of those models of the forming of the early Solar System, which accept higher D/H ratio in comets of the Kuiper Belt than in the Oort Cloud (the former emerged in a colder area of the Solar System than the latter). Moreover, in 2015 the isotopic composition was studied of moisture from Churyumov-Gerasimenko comet (Altwegg et al., 2015) in the same Jupiter family. The moisture’s D/H ratio value for this comet had been found to be equal to (5.3 ± 0.7) × 10−4, i.e., the triple value of the Earth’s ocean. These data indicate that not all comets in the Jupiter family have moisture contents as in the ocean.

Graphic

Figure 1.1 D/H ratio of the present-day ocean a = 155.7 × 10−6 (δDSMOW = 0‰) (Pinti, 2005).

Whereas during the entire time of Earth’s existence the oxygen isotopic composition may have remained unchanged, hydrogen, which more easily loses its lighter isotope, protium, to the outer space may get notably isotopically heavier. The study of oxygen and hydrogen isotopic composition in Isua serpentine of West Greenland in 2012 (Pope et al., 2012) showed that hydrogen in the Archaean ocean water could have indeed been isotopically lighter than in the present-day ocean almost by 25 ± 5‰. Lydia Hallis et al. (2015) from the University of Glasgow also attempted to find moisture of the primeval ocean, which would maintain its original isotopic composition. They believe that they found it in the inclusions of ancient lavas on Baffin Island (Canada), where hydrogen isotopic composition was also notably lighter than in the present-day ocean. In this connection, Hallis et al. proposed a hypothesis that Earth moisture was originally notably lighter than in the present-day ocean and that water had come on Earth directly from the proto-Solar gas and dust nebula, which had formed the Solar System.

Water volume on Earth during the period of its existence could have been notably changing. As major water sources, both the outer space and the subsurface are considered. Moisture from the outer space is coming as a component of the cosmic matter. Approximate estimates (Frank et al., 1986) show that only from numerous smaller comets during the period of Earth’s existence could have come up to (2.2÷8.5) × 10²¹ kg of moisture, which is triple the volume of the present-day ocean. On the other hand, Earth subsurface in the process of matter differentiation by the density gradually loses its most volatile components, among them H2O. These volatile components come to the planet’s surface in the process of volcanic activity and with numerous hydrotherms. Currently, most experts tend to believe that changes in the natural water volume on Earth were mostly associated with degassing of the subsurface, which had begun immediately after the accretion and continued for the entire duration of its history (de Ronde et al., 1994; Kitajima et al., 2001) through volcanoes and hydrotherms.

The atmosphere is a gas shell, which includes about 12.9 thous. km³ of moisture (Table 1.1). The hydrosphere includes natural water on Earth’s surface. Its volume is around 1,338 million km³, which is 96.5% of the identified moisture volume on the planet. Beside liquid water, there is on the surface 24 million km³ of moisture as Arctic and Antarctic ice mountain glaciers. The volume of known ground water is only around 37 million km³, which is around 2.5% of identified water volume. However, a discovery in 2014 by the group of Graham Pearson (Pearson et al., 2014) may notably affect our ideas of the water amount in the planet’s subsurface. The theoretical mineralogy and seismic data suggested for a long time (Jacobsen et al., 2008, 2010; Schmandt et al., 2014) that the major component of the transition zone between Earth’s upper and lower mantle at depths of 410–660 kilometers must be ringwoodite (Mg, Fe²+)2(SiO4), a mineral belonging to the olivine group. However, this mineral was encountered only in meteorites until D. J. Pearson et al. discovered its inclusions, 40 µm in size, within a diamond from Brazil weighing 0.09 g. This discovery confirms that the transition zone between the upper and lower mantle may indeed be composed of ringwoodite. Analyses indicate that up to 1.5% by weight of this mineral is OH− ion, which means that at the boundary between the upper and lower mantle may exist tremendous amounts (up to 1.4·10²¹ kg) of chemically bonded moisture.

Table 1.1 Water amounts on Earth (World water resources …, 2003; Gleick, P. H., 1996).

*Current data after Charette et al., 2010.

There are boundaries to natural water distribution on Earth as liquid water solution. The upper boundary is apparently the altitude of cloud distribution in the atmosphere, i.e., around 13 km over the Earth’s surface. In clouds, water forms suspended droplets, does not have regional limits for its distribution, and is capable of migrating over huge distances at a speed of up to 100 km·h−1 and greater.

The lower boundary of water solutions distribution is deep in the Earth’s subsurface. The portion of the lithosphere which contains mobile water solution is often called hydrolithosphere. The relative ground water content within it declines with depth due to the decrease in the combined porosity, fracturing, channeling and other voids in rocks. Along with it increases the role of moisture emanated by the rocks in the process of epigenesis and metamorphism. This ability of minerals to form moisture is due to the presence in them of either directly H2O molecules (the crystallization moisture) or hydroxyl OH− (the constitution moisture). Minerals containing stoichiometrically bonded hydrogen (first of all in the form of H2O and OH−) are called hydrous minerals (Figure 1.2). With increasing temperature and pressure minerals lose this moisture.

Graphic

Figure 1.2 Hydrogen concentration range in clinopyroxene under various geological environments. The lowermost concentrations are in oxidized rock samples. (Skogby, 1999).

And at last, the minerals include the so called trace hydrogen, which is not part of the mineral structure. The unregulated proton is usually present in the crystalline grid errors (Ohtani, 2007; Demouchy et al., 2016), does not have the stehiometric bond and shows itself up at infrared spectrometry Fourier. It is also a potential source of the moisture and influences rock melting temperature and rheological properties. The minerals where this potential is the main one are called nominally unhydrous minerals. The olivine, pyroxene, garnet, plagioclase and quartz have a capability to accept trace amounts of the structural hydroxile and molecular water in the range of a few ppm to thousands ppm in structural errors of the grid (Hui et al., 2016; Sheng et al., 2016).

In a very high pressure and temperature the mantle rocks may include only hydroxyls with OH− and hydrites with H+. Nevertheless, for the sake of simplicity they are often considered as bonded water. With increasing pressure and temperature, the role of aqueous minerals as potential source of moisture is declining and the role of nominally anhydrous minerals is growing. The moisture of nominally anhydrous minerals is found mostly in the mantle where its reserves may be comparable with the volume of the oceans. It may take part in the water-exchange between the upper mantle and hydro-lithosphere through spreading, subduction zones and hot spots. The moisture of nominally anhydrous minerals, even in negligible amounts, may noticeably affect mechanical strength and rheologic properties of the asthenosphere rocks. Without it, possibly, the plate tectonics would not have occurred (Keppler et al., 2005; Rupke et al., 2013; Demouchy et al., 2016).

With depth, relative content of the pore water and its migratory capability decline. For this reason, the main parameter, which apparently defined the lower boundary of ground water distribution as a liquid water solution, are rock permeability and their capability to provide a continuous or periodic hydraulic connection, which is very important from the water-exchange.

Such boundary may be formed by a zone where the clastic-frictional rock deformation regime is replaced by the quasi-plastic one, i.e., where the rock becomes plastic. Above this zone are maintained porosity, fracturing and faults, which provide through rock permeability not always stable hydraulic connection with the surface. Below this zone even crystalline rocks acquire minimum plasticity and maximum strength. This prevents any hydraulic connection in massifs of underlying rocks. In a plastic rock, moisture is present either as H2O of fluid inclusions or as hydrous oxide OH− in the structures of water-containing minerals or as only hydrogen in the so-called nominally anhydrous minerals.

The zone where plasticity begins is often called the zone of brittle-plastic transition (Scholz, 1988) and forms the last and deepest regional water-confining stratum. When the rocks cool down, this zone separates the cracking front (Monning et al., 2000; Lister, 1974, 1983; Kelley et al., 2002) from the crystallization front (Alt, 1995; Kelley et al., 2002).

The depth to the zone of brittle-plastic transition as water-confining stratum depends mostly on rheologic rock properties and temperature. According to R. Sibson (1986) and Scholtz (1988), its position in acidic quartz-feldspar rocks in the continental crust is defined by temperatures between 300 and 450 °C. Fournier (1991) defined this boundary by temperature of 390 °C (Figure 1.3). In ultramafic peridotites of the oceanic crust a similar zone of brittle-plastic transition is associated with temperature of 700–750°C (Fournier, 2006).

Graphic

Figure 1.3 Schematic concept of positioning the zone of brittle-plastic transition with major geological and seismic features (after Scholtz, 1988).

The main feature of this transition zone is that it is the deepest water-confining stratum whose properties and position in the section depend only on rheologic properties of the rock and temperature (geothermal gradient value). The position of such water-confining stratum has no regional restrictions. In the areas of volcanic activity it approaches the surface, in subduction zones it plunges deeper than 100 km. In the environment of progressive metamorphism the rocks descend through this zone leaving most of their moisture. Under the retrograde metamorphism, on the contrary, rocks ascend through it and are subjected to hydration.

The temperatures within the brittle-plastic transition zone match the conditions of greenstone and epidote-amphibolite facies formation and, as a rule, are higher than critical temperature of water (374 °C). For this reason, with this zone is always associated the process of water solutions’ stratification into high-density mineralized brines and light vaporous fluids. Non-volatile components accumulate in high-density brines directly above the water-confining stratum and volatile components of vaporous fluids move up the section.

All water in the atmosphere, hydrosphere and hydro-lithosphere is associated with the unique geological history of Earth and global water circulation called the hydrological cycle. The oceanic water evaporates, gets into the atmosphere and returns as the atmospheric precipitation. Some of this precipitation penetrates subsurface and becomes ground water. Most of it returns on the surface and again gets into the ocean or atmosphere.

This hydrologic cycle, however, had been notably changing in the course of Earth’s history. It may be assumed that Earth’s atmosphere emerged first, almost simultaneously with Earth’s formation, i.e., 4.51–4.45 billion years ago. The hydrosphere in the form of the ocean over its underlying proto-crust emerged, as it is believed, somewhat later, about 4.0–4.4 billion years ago. The oldest marine deposits are believed to be banded iron ore formations in the stratified sequence of amphibolite and ultramafic rocks, of estimated age of 3.865 ± 11 MMY, on the Island of Akilia in West Greenland (Nutmana et al., 2002; Pinti, 2005). Last emerged the continental crust with its fresh surface water. Its emergence is attributed to Archean, 4.03–3.53 billion years ago. In this order we will review the natural and ground water composition’s formation and properties.

Chapter 2

Atmospheric Water

The atmosphere contains around 12.9–13.8 thous. km³ of moisture, i.e., only 0.001% of its total amount on Earth. Had all this moisture precipitated on Earth’s surface it would have formed a layer only 25–27 mm thick. Despite its negligible amount, it plays a huge role in the distribution of mineral components and fresh water on Earth. The process of its formation is similar to the operation of a giant water distillation unit, which separates mineral components and H2O. The atmosphere is in fact disrupting the circulation of mineral components. The amount of salts, which the ocean is losing through the atmosphere (around 0.6·10⁹ t·year−1) is negligible. That is the reason why natural water, which precipitates from the atmosphere is the freshest on Earth. It is exactly for this reason that the formation of surface and ground water composition should be reviewed starting with the atmospheric precipitation.

As mentioned earlier, the atmosphere could have emerged prior to the ocean; therefore it could have participated in its formation. It is believed that at the earliest stages in the history of our planet (in Katarchean, 4.56–4.0 billion years) the entire water on the surface of Earth was atmospheric and its major components were CH4, NH3, CO2, H2 and noble gases. At that time the atmosphere almost did not allow the Solar light through and the pressure and temperature on the planet’s surface were very high. Only in Archean (4.0 – 2.5 BY) did the atmospheric water reach the surface and form the ocean (Pinti, 2005), which absorbed a substantial part of well soluble NH3 and CO2. At the same time, relative concentration of volatile N2 and CO2 in the atmosphere increased. Later, due to the propagation of live organisms and formation of carbonate deposits CO2 almost totally disappeared, and the atmosphere became transparent. Approximately 3.5 BY ago appeared photosynthesizing micro-organisms (cyanobacteria), which formed O2 and thereby increased Eh of the atmosphere. The cyanobacteria, having excess of CO2, fast multiplied and saturated with oxygen not only the atmosphere but also the ocean. The emergence and propagation of aerobic organisms facilitated establishment of some oxygen equilibrium. It provided for the stability of O2 concentration both in the atmosphere and in the ocean. Already in Proterozoic the main components of the atmosphere became N2, O2 and Ar.

Volume of the present-day atmosphere includes 78% N2, 21% O2 and 0.9% Ar. Its maxim moisture concentration depends on temperature and regularly decreases from 30 g·m−3 at 30 °C to 8 g·m−3 at 10 °C and 0.3 g·m−3 at –10 °C. For this reason, maximum air humidity changes from almost 2.6% near the equator to 0.2% at the latitude of around 70°. Most moisture is concentrated in the lower strata of the atmosphere (up to 70% below 3.5 km and around 90% below 5 km).

Within a year through the atmosphere rotates around 520–577 thous. km³ of moisture. Most of it (82.1%) comes from the ocean surface. Thus, atmospheric water is replaced during a year around 45 times, i.e., every eight days.

Water droplets form when maximum humidity is reached. This is facilitated by lowering of the temperature and by the presence of dust particles. Accumulations of these droplets or tiny ice crystals (4 to 140 µm) form clouds. While dropping, rain droplets merge and increase in size to 0.5–3.0 mm. The layer of atmospheric precipitation dropping within a year varies depending on the climate between 50 mm in deserts and 3,000 mm in humid rain forests.

Isotopic composition of the atmospheric moisture is controlled mostly by the processes of its evaporation and condensation. At evaporation, isotopically lighter H2O molecules quicker and in greater numbers pass to vapor. As a result, gaseous moisture under normal conditions becomes almost poorer by 8% in deuterium and by 0.9%, in ¹⁸O than in the source water. Thus, isotopic composition of the atmospheric moisture is always somewhat deficient in heavy isotopes compared to the source. Average vapor isotopic composition in the ocean has δD around –22‰ and δ¹⁸O around –4‰. On the contrary, at condensation rain droplets enrich in heavy isotopes and become isotopically heavier than the source vaporous moisture. Therewith moisture preserved in the atmosphere after the rain becomes isotopically even lighter. This isotopic fractionation defines the distribution of stable isotopes (δ¹⁸O and δD) in the atmospheric precipitation. Moreover, the coefficient of such fractionation increases with the growth of temperature. Thus, the δD and δ¹⁸O values turn out tied with one another by an equation called the global meteoric water line:

When the air mass moves from the equator to the poles and is cooled down, it loses moisture. Thus, moisture preserved in the atmosphere becomes isotopically lighter. For this reason, the isotopic composition of meteoric water varies between values close to 0‰ for δ¹⁸O and 10‰ for δD at the equator and –20‰ for δ¹⁸O and –200‰ for δD, in polar areas (Figure 2.1). In Greenland, δD values are within the range of –210‰ to –310‰ and δ¹⁸O values, between –36,6 and –13,6‰ (Craig, H. et al., 1965; Bonne et al., 2014).

Graphic

Figure 2.1 Contours of δ¹⁸O value in January precipitation relative SMOW standard. The data are received from IAEA (Lawrence et al., 1991).

As most moisture comes from the side of the oceans, its δD and δ¹⁸O values decline into the heart of continents (approximately by 0.7% deuterium per 100 km). The atmosphere of continents loses its remaining moisture usually when encountering the mountains. For this reason most isotopically light meteoric water precipitates closer to the mountain tops. V. M. Mukhachev (1975) believes that deuterium unfavorably affects all living creatures, and the isotopically light moisture is identified with the life-giving water. If so, the greatest amount of isotopically light life-giving water is hidden in the ice of the Antarctic, Greenland and Arctic.

Besides, in any limited territory moisture isotope composition experiences local fluctuations as affected by the local climatic conditions (Figure 2.2).

Graphic

Figure 2.2 Seasonal variation in the precipitation value of δ¹⁸O and δ²H. At the top, local precipitation observed at the Tien River, Vietnam. (Nguyen et al., 2012). At the bottom, the daily precipitation observed at Yakutsk (Sugimoto et al., 2012).

The mineral composition of atmospheric precipitation begins to form in the ocean. Salts and moisture are getting in the atmosphere separately. The salt is present in the atmosphere as small suspended particles. These particles are the dry residue from small droplets of splashes and are composed mostly of NaCl. Onland, solid particles in the atmosphere are either the dust raised by wind or ashes from volcanic eruptions. There, they are composed of CaCO3, MgCO3, CaSO4·2H2O, Na2SO4, MgSO4, aluminum silicates, organic matter and even live microorganisms. These particles form aerosols, which may be carried by the atmosphere over substantial distances. The aerosol concentration rapidly declines with increasing altitude, and the composition of the atmospheric downfall forms mostly at the moment of its precipitation. For this reason, the precipitation’s mineralization is, as a rule, in inverse relation with their amount. For instance, in Sankt-Petersburg area 10 mm of the precipitation has mineralization of 11 mg·l−1 and 20 mm, 6 mg·l−1. Rains in fact are cleaning up the atmosphere by removing solid suspended matter from it. Only in conditions of very high surface temperature some rain moisture evaporates while dropping, and mineralization of its droplets increases. On rare occasions the rain may even not reach Earth’s surface. Then the so-called dry rain or virga is observed.

The atmospheric precipitation is distinct among the other natural water by the lowermost mineralization, 3–4 to 50–60 mg·l−1, and is very rarely greater than 100 mg·l−1. The mineralization values depend mostly on climatic conditions and increase with the decline of the precipitation and growth in atmospheric dustiness. The lowermost mineralization is observed in polar areas. In the Antarctic, the precipitation mineralization usually does not exceed 3–4 mg·l−1. In Russian plain territories, average mineralization of the meteoric water increases from the north to the south according to landscape-vegetation zones from 10 – 15 mg·l−1 in the northern coastal tundra to 20 mg·l−1 in the forest zone and to 60 mg·l−1 in steppes and forest-steppes. As a rule, the precipitation mineralization notably declines up the mountain slopes. Maximum contents of mineral components, sometimes greater than 200 mg·l−1, is observed in dry steppe and desert areas. Rarely, in arid areas and above large industrial cities such mineralization reaches 270–550 mg·l−1.

A. A. Kolodyazhnaya (1963) estimated that in European Russia between 50 and 80 ton of salt annually dropped per each square kilometer of the surface, and in some U.S. and British cities, up to 240–360 ton.

Precipitation’s macro-component composition depends on the conditions of their formation. In the ocean and marine territories they are dominated by Cl− and Na+. Moving gradually inland, the fraction of these components notably declines. Whereas in the oceanic atmospheric precipitation Cl− content may reach 100 mg·l−1, in the coastal zone it declines to 10 mg·l−1 and exponentially declines with the distance from shore. At a distance 20 km content of Cl− is halved and at a distance of 100 km drops to 1 mg·l−1. In Europe, the chlorine content declines from 10–15 mg·l−1 in the coastal zone to 1–2 mg·l−1 at a distance of 200–400 km. The same pattern is observed in the Antarctic. As a result, Cl− is gradually substituted with SO4²− and carbonates, and Na+, with Ca²+ (Figure 2.2). For this reason in the Russian territory the precipitation dominated by Cl− and Na+ is observed only in coastal areas of the Arctic, the Black Sea region and the Far East (Figures 2.3, 2.4).

Graphic

Figure 2.3 Annual average SO4²– and Cl− mineralization and concentration in the atmospheric precipitation of Eastern Europe, mg·l−1 (Drozdova et al., 1964).

Graphic

Figure 2.4 Annual average mineralization of the atmospheric precipitation, mg·g−1 (a) and salts annually coming with it, ton·km−2 (b) in Eastern Europe (Nikanorov, 2001).

Table 2.1 Composition of atmospheric precipitation (mg·l−1) for 1958–1961 (Nikanorov, 2001; Shvartsev, S. L., 1998; Hydrochemistry, 2001).

In their salt composition the atmospheric precipitation usually belongs to hydrocarbonate-sodium and sulphate-sodium type (after Sulin) because in them rNa > rCl1. Precipitation with rNa < rCl is very rare (in conditions of arid climate, in particular, in Baghdad - Al-Aili et al., 2007).

As chlorine is capable of volatilizing, relative content of macro-components in the atmospheric precipitation is usually compared with Na+ concentration. For such comparison is used the fractionation factor, which is determined from the following equation (Appelo et al., 1994):

(2.1)

where Fi is fractionation factor, Ratm. and Roc. are concentration ratios of component i and Na+ in the atmospheric precipitation and in the ocean, respectively.

Moving inland of continents, rCl fractionation factor usually drops below 1, whereas for the other components it notably increases. Maximum growth, up to hundreds, is observed for the factor of carbonate alkalinity and to a smaller extent for Ca²+ and SO4²–. Apparently, these components are coming from the surface of continents. Their contribution is in direct correlation with air dustiness and in inverse correlation with the amount of the precipitation. Air emissions of the industrial waste also facilitate contamination of the atmospheric precipitation. In particular, sulfur dioxide (SO2), which forms in great amounts at burning of fossil fuels, reacts with the atmospheric oxygen and moisture and increases sulfate content:

Annually, up to 0.5–0.6 ton, sometimes even 1.0–1.5 ton of sulfur entered per every square meter of European Russia.

In the mountain areas of the continents, in conditions of very low dustiness the anions are dominated by HCO3−. Such precipitation occurs at high altitudes, in the Caspian and Black Sea coastal areas of the Caucasus. Hydrocarbonates are also dominant in the precipitation in the Antarctic.

It deserves attention that there is a relatively high content of K+ in the precipitation. As a result, values of the Na+/K+ ratio in them sometimes decline to 1.5–2.0, which is almost by an order of magnitude lower than in most natural water.

Nitrogen compounds in the precipitation are ammonium NH4+ (0.2–1.2 mg·l−1) and nitrate NO3− (0.03–2.3 mg·l−1). Nitrates form by oxidation of the atmospheric molecular nitrogen at the moment of electric discharge (average content of nitrates in the precipitation in the absence of thunderstorms is seven times lower than in thunderstorms) and also at oxidation of nitrogen compounds brought in with industrial waste and volcanic eruptions. Between 0.2 and 0.4 ton of bonded nitrogen per square kilometer of the surface in European Russia comes with the precipitation.

It appears that all elements of Mendeleyev Table may be identified in each drop of rain water. By measured concentrations the micro-components, according to V. M. Drozdova et al. (1964), may be subdivided into three basic groups: with the concentrations greater than 5·10−5 mg·l−1 (Si, Fe, Mn, P, Cu, Zn); with the concentrations between 5·10−5 and 5·10−6 mg·l−1 (Li, Sr, V, Be, B, Ba) and with the concentrations less than 10−6 mg·l−1 (Sn, Bi, Pb, Ag, Cr, Co, Mo, Ti). Notable increases in the concentrations of some micro-components are associated with technogenic contamination of the atmosphere. For instance, mercury concentration in the precipitation over some cities may reach 2.7·10−4 – 14·10−4 mg·l−1. Lead application for lowering gasoline detonation brings it in the atmosphere with the resulting increase of its content sometimes by 20 times.

Especial danger is associated with technogenic radioactive components. Their natural concentrations are harmless. However, due to surface nuclear explosions and emergencies at the nuclear power generating stations these concentrations notably increased. Maximum nuclear testing occurred in 1954–1958 and in 1961–1962. Between 1963 and 1980 surface tests were conducted only by France and China. In the Chernobyl accident in April of 1986 the height of the first ejection reached 1,200 m and thereafter did not exceed 400 m. In the process, the maximum surface contamination was limited by the radius of 60 km. Some radioactive dust reached the stratosphere (the altitude of 10–50 km) and was spreading all over Earth surface for many months. The radioactive precipitation contained a few hundred different radionuclides. The greatest threat for the population comes from radioactive isotopes ¹⁴C, ⁹⁰Sr, ⁹⁵Zr and ¹³⁷Cs.

Gaseous components enter the meteoric water directly from the atmosphere whose both components’ isotopic composition are stable in time. This enables the utilization of concentrations of some atmospheric components or their ratios, in particular for the noble gases, as constants.

As raindrops are born in the atmosphere they, obviously, are maximum saturated with its gas components. For this reason the gas composition of rain is defined by the composition of ambient air and value of the atmospheric pressure. Therefore, with declining atmospheric pressure up the atmosphere also declines the content of dissolved gas components.

At the pressure of 1 atm in the fresh water of the precipitation can dissolve up to 12 mg·l−1 of molecular nitrogen and up to 8 mg·l−1 of oxygen. Among the rest of gas components special interest cause CO2, CH4 and noble gases.

The current estimate of CO2 concentration in the atmosphere is a little over 0.040%, although 150 years ago it was approximately 0.026%. Due to a very high solubility the content of CO2 in rainwater reaches ~ 0.4 ml·l−1. Therefore, the pH value of the meteoric precipitation should be 5.6. Indeed, the pH value of rainwater is usually less than 7.0 and often is within the range of 5 to 6. Alkaline rains are relatively rare and are associated with the presence of alkaline dust. Very low pH values are usually due to the atmosphere contamination with sulfur dioxide SO2 or nitrogen oxide NOx. In these cases the pH value can fall below 4.

Concentration of CO2 in the atmosphere deserves attention because it retains the short-wave heat reflection from Earth, thereby creating the greenhouse effect. In the absence of this effect average temperature on Earth would not exceed minus 15 °C. Doubling in the CO2 content because of burning fossil fuels and forest fires can cause increase in global temperature by 2–3 °C. Some scientists believe that because of this the ocean level in the last 100 years has risen by 15 cm.

Negligible amounts of methane are present in the air (~1.75·10−4% by volume) (Wuebbles et al., 2002). Long-term observations indicated that its content over the industrial period more than doubled. Isotopic composition of methane’s carbon in the atmosphere is described by value of δ¹³C around – 47‰ (Khalil, 2013). Methane, like carbon dioxide, absorbs infrared radiation and also facilitates the emergence of the greenhouse effect. However, of much greater interest is a suggestion by D. Blake and S. Rowland (1988) that methane in the stratosphere decomposes and facilitates formation there of small ice crystals, in the presence of which chlorine acquires the capability to destroy the ozone layer. Blake and Rowland suggest that in the last 40 years moisture content in the stratosphere could have increased by 28%, and in the last 200 years, by 45%.

Chemical hydrologists are also interested in argon and other noble gases whose content and isotopic composition are very stable. The content of argon in the atmosphere reaches 0.934%. For this reason its content in rainwater should not exceed ~0.4 ml·l−1. In Earth’s atmosphere it is represented by stable isotopes: ³⁶Ar (0.337%), ³⁸Ar (0.063%) and ⁴⁰Ar (99.600%). The isotopic composition of the atmospheric argon is described by a constant ratio ⁴⁰Ar/³⁶Ar equal to 295.6 and is used as the isotopic standard. Helium content in the atmosphere is negligible (0.0005239%), and its concentration in rainwater should be close to 1.0·10−4 ml·l−1. Helium is composed of two stable isotopes, ³He and ⁴He, with the ratio of 1.39·10−6 (Mamyrin et al., 1970). The other noble gases are studied relatively rarely.

Of great interest is the cause of radioactive cosmogenic isotopes in the atmosphere, especially ¹⁴C and ³H whose concentrations are relatively constant due to the stationary equilibrium between the rates of their formation and decay.

Organic matter enters the atmosphere with spores, pollen, live organisms or directly with the dust as detritus. The live organisms are mostly aerobic and autotrophic microorganisms. Organic matter in the atmospheric precipitation includes fat acids, hydrocarbons, ethers, amino acids and other compounds. Nevertheless, the content of organic matter in the atmospheric precipitation is negligible. Concentration of organic carbon Corg. varies between 1.7 and 3.4 mg·l−1. Annually, up to 0.4 t of organic matter per a square kilometer reaches the surface in the European Russia with the atmospheric precipitation.

Thus, meteoric water is an important link of the hydrologic circulation where H2O is separated from salts and transferred to the territory of continents. Its composition is defined mostly by climatic and geographic factors. A similar effect of geological factors is limited, has a different nature in the ocean and on the continents, and is believed to be associated mostly with volcanic activity.

Summarizing, the main distinctive features in the composition of atmospheric precipitation are:

Their isotopic composition is controlled by the law of the Global Meteoric Water Line and becomes lighter from the equator to the poles of Earth;

Their mineralization is lowermost (less than 100 mg·l−1) among natural waters of Earth;

As a rule, in the atmospheric precipitation rNa > rCl, and they belong to the hydrocarbon-sodium or sulphate-sodium type (after Sulin).

The content of gas components in the atmospheric precipitation is controlled by the composition of the atmosphere.

The atmospheric precipitation is saturated with O2 and for this reason have relatively high oxidation-reduction potential, between 0.35 and 0.70 v.

The atmospheric precipitation contains sufficient concentration of CO2 to make the pH usually lower than 7.

1 The letter r before the ion symbol means that its concentration used in meq·l−1.

Chapter 3

The Oceanic Crust

Oceanic crust covers the area of 306 million km², i.e., 60% of the surface of our planet. It is practically completely covered with the oceanic water. The ocean contains 95.6% of the entire water on Earth.

Two stages are identified in the topography of the present-day oceanic floor. The upper one is positioned at depths shallower than 200 m and is called shelf. It covers 7.5% of the ocean floor area. Most of the shelf belongs to the seas. The lower stage (at depths of 4.0 to 5.5 km covers 77.5% of the entire oceanic area and forms abyssal plains. To this stage are also attributed depressions deeper than 4.5 km called deep cavities. These stages are separated by the continental slope at depths of 200 to 4,000 m. The shelf and continental slope belong to the continental crust and the abyssal plains with deep-water depressions, to the oceanic crust.

Thus, water exchange between the oceanic crust and the atmosphere is possible only through water stratum of the world ocean, nearly 4.5 km thick.

3.1 The Ocean Waters

Notably, the ocean and its seas are a major source of fresh water in the atmosphere and on continents. Indeed, the ocean and its seas are the medium where sedimentary rocks form. Their water, which penetrated subsurface with sedimentary rocks is called sedimentogenic water. At the same time the ocean and its seas are closely bonded with underground water of the oceanic crust and atmosphere through continuous water exchange, which actively affects the formation of their composition.

3.1.1 Oceanic Water Regime

The world ocean (Pacific, Atlantic, Indian and Arctic) together with seas covers 71% of Earth’s surface and contains 1,332.48 million km³ of water (Charette et al., 2010). Seas are near-shore parts of the ocean protruding into the continent.

Its water composition forms with the active participation first of all of exogenous factors, solar radiation (which defines temperature regime of the ocean surface), Lunar and Solar gravity (which causes high and low tides) and water exchange with atmosphere and continents. All these factors cause the stratification of ocean water by the density, hence by the composition.

Currently, the ocean annually loses to the atmosphere 505 thous. km³ of moisture (water layer 1.40 m thick). Most of this moisture (405 thous. km³·year−1) returns back, and only 100 thous. km³·year−1 gets off the continent and forms the surface runoff. The ocean loses from its surface only around 1.8–1.9 ton·year−1 of salt (mostly NaCl) (Edelstein, 2005). Most of it returns back in the ocean and only ~30% gets on the present day continents.

Due to the solar radiation and atmospheric precipitation, near the surface forms relatively thin but stable water stratum with quite variable properties and composition. With it are associated maximum temperature range, –2 to +22 °C, and water layer most illuminated by the sun (so-called euphotic or photic zone). The lower boundary of this zone is defined by the depth, which 1% of the sunlight reaches. The euphotic zone thickness in the Sargasso Sea transparent water reaches 150–200 m, in temperate latitudes it is around 40 m. Its average thickness is around 80 m. This is the stratum of most favorable conditions for photosynthesis and habitation of flora and fauna. The euphotic zone of the ocean is a major source (up to 90%) of oxygen in Earth’s atmosphere.

High-density cold water in circumpolar ocean areas submerges and is positioned at the ocean floor. Its composition is highly stable, and the temperature (unless affected by hydrotherms) is within a narrow range of –2 to +4 °C. For this reason everywhere except the highest latitude at the surface is positioned warm water with variable temperature and composition and at the ocean floor, cold water of relatively stable composition and low temperature (Figure 3.1). They are separated by the water stratum of drastically changing with depth properties and composition, so-called pycnocline. This stratum with drastic change in water density is almost coincident with the thermocline, i.e., the interval of temperature decline and stabilization. The thermocline depth in tropical and equatorial zones reaches 300–500 m and decreases toward high latitudes.

Graphic

Figure 3.1 Laminated structure of the ocean and relative volume of each zone (schematic diagram). Note that the «deep» zone next to the poles approaches the surface (after Driver, 1985).

The near-bottom cold water migrates from polar areas to the locations where it periodically discharges on the surface in the form of upwelling. The discharge of near-bottom water on the surface is usually observed near the western shores of continents where continuous winds drive the surface warm water off and allow a cold near ocean floor water to come to the surface (Figure 3.2). Minor upwelling sometimes arises near the western shores of the Crimea and Caucasus.

Graphic

Figure 3.2 Five most productive zones of near-shore upwelling (dotted areas) and of the atmospheric pressure system at sea level (anticyclones) affecting upwelling. Arrows show the approximate positions of major currents (U. Odum, 1986, p. 292).

With the emergence of the continental crust appeared a new influence factor, the continental runoff, which brings in the ocean fresh water mass with average mineralization of 0.78 mg·l−1 (Edelstein, 2005). The value and composition of this runoff depends on the areal extent, elevation above the sea level and climatic conditions on the continents. Currently, its value is estimated at about 37.3 thous. km³·year−1 (Dai et al., 2002) to 42 thous. km³·year−1 (Edelstein, 2005). Along with water, this runoff introduces in the ocean mineral matter, both suspended and dissolved. Average total amount and composition of this runoff into the World ocean from rivers depends on the area, elevation over the sea level and climatic environment of the continents. Currently its value is estimated at 37.3 (Dai et al., 2002) to 42 thous. km³·year−1.

Average duration of water stay in the world ocean may be estimated using the ratio of its volume and the average value of annual rate at evaporation: 1,338,000 thous. km³/ 505 thous. km³/year) = 2,650 years. Therefore, during Holocene, water in the world ocean could have been replaced only four times.

Radiocarbon analyses also enable estimating the age of the oceanic water. According to Matsumoto (2007) and Gebbie et al. (2012), this age varies within a broad range. At depths over 1,500 m the youngest water is found in the circumpolar areas and the oldest (up to 1,000–1,100 years), in the northern Pacific. It needs to be remembered, however, that these age estimates may be undervalued due to underestimation of ¹⁴C removed together with carbonate deposits.

3.1.2 The Oceanic Water Composition

The oceanic water composition is very stable as a result of centuries-long mixing and a stable balance of dissolved components. The oceanic water composition is described by isotopic composition of H2O and contents of mineral and gas components. The distribution of biogenic elements is strongly affected by biochemical processes.

The isotopic composition of the oceanic water, especially deeper than 500 m, is very stable. Harmon Craig (1926–2003) used it to propose in 1961 the oceanic water at great depths as the standard of hydrogen and oxygen isotopic composition. Three strata may be identified by the environment of H2O isotopes’ distribution in the oceanic water: the surface one (to 500 m deep), transitional and deep (1,000 m and deeper).

H2O isotopic composition in the upper stratum strongly depends on isotopic fractionation associated with evaporation and condensation. Evaporation is removing light isotopes and making the ocean water isotopically heavier, whereas the atmospheric precipitation is mixing with it and making it isotopically lighter. This isotopic lightening is made especially strong by addition of the melt-water from polar ice and icebergs. Major providers of isotopically light moisture are the Antarctic and Greenland. For this reason, water isotopic composition in the surface stratum varies within the ranges of around 35‰ for δD and around 3‰ for ¹⁸O (Ferronsky, 2015). Overall, the oceanic moisture of the oceanic surface (above thermocline) experiences a regular isotopic heaving from the polar areas to the equator. In the deep ocean strata H2O isotopic composition is very stable with δD variations not exceeding 4‰, and ¹⁸O variations, not exceeding 0.3‰ (Ferronsky, 2015). For the bottom water with temperature around 3 °C H2O isotopic composition is described by δD values between –0.14 and –0.6‰ and δ¹⁸O values between –0.16 and –0.17‰ (Bohlke et al., 1994; Reeve et al., 2011).

The mineral composition of the oceanic water, despite a huge volume and distribution area, is surprisingly constant, thus reflecting the stability of its water and salt balance. There are reasons to believe that the ocean volume and composition varied only insignificantly, at least since Paleozoic.

Average mineralization of the oceanic water currently is 35 g·l−1. The permanency of sea water mineralization was noted by R. Boyle (1627–1691) in the seventeenth century. The sea water composition was first determined in 1859–1865 by Johann Georg Forchhammer (1794–1865). Currently, in this composition are discovered almost all elements of D. I. Mendeleyev’s Table. The anions are dominated by chlorine; second in value are sulphates. The main cations are Na+ and Mg²+. The oceanic water is quite hard and belongs to the chloride-magnesium type in salt classification by V. A. Sulin. Mineral composition of the present-day ocean is shown in Table 3.1.

Table 3.1 Major ions of the oceanic water after S.V. Bruyevich (1966) (per 1 kg at S=35.00‰ and Cl=19.375‰).

Even at significant mineralization changes the ratios of major components remain practically constant, indicating a good mixing of the sea and ocean water. That allows the use of these ratios (Br/Cl=0.0033; SO4/Cl = 0.14; Na/Cl=0.55; Mg/Cl=0.066, etc.) as the most reliable criterion of the presence of oceanic water. Stability of these ratios was observed by William Dittmer (1833–1892) based on 77 analyses of samples collected in different parts of the ocean. In 1884 this was defined as a law, which currently bears his name. Dittmer’s law maintains that in water of the open ocean regardless of absolute concentrations quantitative ratios of major components in its basic salt composition are