FeSS S2O2aq FeS2S SOeqn 616

The sulphite (SO2-) is subsequently oxidized to SO4-. Sedimentary pyrite, formed as a byproduct of sulphate reduction in marine sediments, is a major sink for seawater SO4-. The presence of pyrite in ancient marine sediments shows that SO42- reduction has occurred for hundreds of millions of years. On a geological timescale, removal of SO4- from seawater by sedimentary pyrite formation is thought to be about equal to that removed by evaporite deposition (Section 6.4.2). Compilations of pyrite abundance and accumulation rates are used to calculate modern SO4- removal by this mechanism and to derive the estimate in Table 6.2.

Sulphate reduction (eqn. 6.12) also produces HCO-, and this anion slowly diffuses out of the sediment into seawater, accounting for about 7% of the HCO-flux to the oceans. The slow diffusion also means that HCO- may build up to sufficiently high concentrations in sediment pore waters for the ion activity product of Ca2+ (from seawater) and HCO3- to exceed the solubility product for

Seawater / Pillow basalts

1

r"f ^^

^^ Emplacement of dolerite ——

•p

~ 2

* T

op c

magma nber

Fig. 6.13 Simplified structure of a fast spreading mid-ocean ridge (e.g. East Pacific Rise).

Fig. 6.13 Simplified structure of a fast spreading mid-ocean ridge (e.g. East Pacific Rise).

CaCO3 (eqn. 6.6). This allows CaCO3 to precipitate as nodules (concretions) in the sediment. These are not quantified in the global budget as they are volu-metrically small sinks of CaCO3.

6.4.7 Hydrothermal processes

The hydrothermal (hot water) cycling of seawater at mid-ocean ridges has a profound effect on the chemistry and budgets of some major and trace elements in seawater. To understand the chemical changes it is necessary to know a little about geological processes at mid-ocean ridges.

Basaltic ocean crust is emplaced at mid-ocean ridges by crystallization from magma that is sourced from a magma chamber at shallow depth (about 2 km) below the ridge. The magma chamber and newly emplaced crust is a discrete heat source, localized below the ridge (Fig. 6.13). Successive emplacement of new ocean crust gradually pushes the older crust laterally away from the ridge axis at rates of a few millimetres per year. This ageing crust cools and subsides as it travels away from the ridge axis. The resulting thermal structure, i.e. a localized heat source underlying the ridge with cooler flanking areas, encourages seawater to convect through fractures and fissures in the crust. This convection is vigorous close to the ridge axis but more passive on the off-axis flanks (Fig. 6.14).

The deep waters of the oceans are cool (around 2-4°C) and dense relative to overlying seawater. This dense water percolates into fissures in the basaltic crust. Slowly it penetrates deep into the crust, gradually heating up, particularly as it approaches the heat source of the underlying magma chamber. This massive heat

Fig. 6.14 Convection systems at mid-ocean ridges. Active circulation at the ridge axis is driven by magmatic heat. Off-axis circulation is driven by passive cooling of the crust and lithosphere. Modifed from Lister (1982).

source warms the water, causing it to expand and become less dense, forcing it upward again through the crust in a huge convection cell (Fig. 6.15). We can view these convection cells as having a recharge zone and low-temperature 'limb' of subsiding seawater, a hot reaction zone closest to the magma chamber, and a high-temperature rising 'limb' of chemically modified seawater discharging at the seabed (Fig. 6.15). The overall process is called 'hydrothermal' (hot water) convection. It is currently estimated that about 3 X 1013kgyr-1 of seawater is cycled through the mid-ocean ridges of the Earth's crust. It therefore takes about 3.3 X 107 years to cycle the entire volume of seawater through the axial part of the mid-ocean ridges (Table 6.5).

It is not possible to measure directly the maximum temperature to which water becomes heated in the basaltic crust. However, hot springs of hydrothermal water discharge from the seabed at the apex of the convection cell. Temperature measurements taken from ridge axis hot springs range from 200 to 400°C (average around 350°C). This implies that temperatures in the hot reaction zone above the magma chamber (Fig. 6.15) are typically not less than 350-400°C. Owing to chemical reactions between this convecting hot water and the basaltic crust (shaded region on Fig. 6.15), these waters are acidified (typical pH 5-7) and rich in dissolved transition metals leached from the crust. Iron (Fe), manganese (Mn), lead (Pb), zinc (Zn), copper (Cu) and hydrogen sulphides rapidly precipitate a

Hot springs and black smokers

Hot springs and black smokers

Fig. 6.15 Hydrothermal convection at a mid-ocean ridge.
Table 6.5 Estimates of ocean volume circulation times through hydrothermal systems. Modified from Kado et al. (1995).

Axial hydrothermal convection

Off-axis convection

Recycling through

(350°C fluid)

(20°C fluid)

'black smoke' plume*

3.3 X 107yr

5.5 X 105yr

2.8 x 103yr water

2.4 x 105yr reactive

elements

* Extrapolated from study of the Endevor Ridge. Although the ocean volume is cycled through the plume in 2.8 X 103 years, the rate of reaction of the plume with seawater is much slower. To strip the reactive elements from seawater it is necessary to cycle the ocean volume approximately 100 times through the plume (2.4 X 105yr).

* Extrapolated from study of the Endevor Ridge. Although the ocean volume is cycled through the plume in 2.8 X 103 years, the rate of reaction of the plume with seawater is much slower. To strip the reactive elements from seawater it is necessary to cycle the ocean volume approximately 100 times through the plume (2.4 X 105yr).

cloud of iron, zinc, lead and copper sulphides and iron oxides on injection into cold, oxic oceanic bottom waters. This sulphidic plume of particles identifies clearly the location of the hot springs and gives rise to their colloquial name — 'black smokers' (Plate 6.1, facing p. 138).

The high temperatures encountered in hydrothermal circulation cells at mid-ocean ridges increase substantially the rate and extent of chemical reaction

Fig. 6.16 Location of known mid-ocean ridge hydrothermal systems in the oceans. After Baker et al. (1995).

between seawater and basaltic ocean crust. However, estimates of the element fluxes involved in these processes are uncertain, mainly because representative sampling in these remote environments is difficult and expensive. The flux estimates are based on a few studies at individual sites on the East Pacific and Mid-Atlantic ridges (Fig. 6.16). Global fluxes have been calculated from these sites, using various geophysical and geochemical approaches. A major problem however, still to be resolved, is a rigorous quantification of the amount of hydrothermal activity occurring at high temperatures near the ridge axis, versus lower temperature circulation on the ridge flanks (Fig. 6.14 & Table 6.5). This is important, because temperature affects the degree, rate and even direction of some chemical reactions. Despite these problems, for some elements the direction of the fluxes agree from site to site allowing construction of tentative global fluxes (Table 6.2). We should, however, note that the magnitudes of the fluxes in Table 6.2 are uncertain.

In the following section we describe the effects of hydrothermal activity on major ions in the oceans: the effects on minor components in the oceans are discussed in Sections 6.5.2 and 6.5.5.

Hydrothermal reactions as major ion sinks

Of the major elements, the case for magnesium removal from seawater during hydrothermal cycling at mid-ocean ridges is most convincing. Experimental work and data from many black smokers (Fig. 6.16) suggests that the hydrothermal fluids exiting from the crust have essentially zero magnesium concentration. This implies that magnesium is removed from seawater by reaction with basalt at high temperature. The precise chemistry of this reaction is not known, but it can be represented as the generalized reaction:

(fayalite) (seawater)

^ Mg2Si3O6 (OH) 4( S) + 7Fe3O4( s) + FeS2(S) + 8SiO2(aq)

(sepiolite) (magnetite) (pyrite) (silica) eqn. 6.17

The basalt, represented here by iron-rich olivine (fayalite), is leached of its iron and hydrated by seawater, whilst Mg2+ from seawater is used to form the magnesium clay mineral (sepiolite in eqn. 6.17), which represents altered basalt. The formation of magnesium clay mineral also removes OH- from water to make the (OH)4 component of the clay mineral in equation 6.17. In laboratory experiments, it is this removal of OH- from H2O that leaves the fluid enriched in H+, explaining the acidity of hydrothermal fluids. The H+ does not show up in the products of equation 6.17 because the equation is a summary of a number of processes going on over time (see Section 2.4). However, acidity generation is an important feature that, along with complexation by Cl- anions (Box 6.4), enhances iron solubility. The reaction (eqn. 6.17) also predicts the formation of iron oxide (Fe3O4), iron sulphide (FeS2) and silica, all of which are found at hydrothermal vent sites. Although it is difficult to quantify the amount of Mg2+ removed from seawater by this process, it is probably the most important Mg2+ sink in the modern ocean. There is also uncertainty about the fate of Mg2+ in altered basalt (sepiolite in eqn. 6.17) as it moves away from the ridge axis during seafloor spreading. There is evidence that Mg2+ is leached from altered basalt by cold seawater. If large amounts of Mg2+ are resupplied to seawater by such low-temperature basalt-seawater interactions, then mid-ocean ridge processes may not cause net Mg2+ removal from seawater.

Sodium is by far the most abundant cation in hydrothermal fluids, simply because the fluid is sourced from seawater (Table 6.1). It has long been thought that Na+ must be removed from seawater at mid-ocean ridges, mainly because the global Na+ budget does not otherwise balance. The existence of Na+-enriched basalts (spilites), believed to have formed by reaction with seawater at high temperature, is tangible evidence that Na+ removal from seawater occurs during mid-ocean ridge hydrothermal activity. The formation of sodium-feldspar (albite) probably accounts for the removal process from the fluid to the altered basalt, but data to quantify fluxes are scant. Hydrothermal activity probably accounts for no more than 20% of the total Na+ removal from seawater, which

Table 6.6 Changes in seawater major constituents upon reacting with mid-ocean ridge basalt at high temperature. Hydrothermal fluid data are typical ranges from Von Damm (1995).

Seawater*

Hydrothermal fluids

A

Constituent

(mmol l-1)

(mmoll-1)

(mmoll-1)

Mg2+

53

0

-53

Ca2+

10

10-100

0-90

K+

10

15-60

5-50

SO4-

28

0-0.6

-28

H4SiO4

0.1

5-23

5-23

*Data from Table 6.1.

A, difference between typical range in hydrothermal water and seawater.

*Data from Table 6.1.

A, difference between typical range in hydrothermal water and seawater.

on a geological timescale is dominated by the formation of evaporites (Section 6.4.2).

Hydrothermal reactions as major ion sources

The chemistry of hydrothermal fluids indicates that basalt-seawater interactions are a source of some elements that have been stripped from ocean crust and injected into seawater. Data from hydrothermal fluids show that both Ca2+ and dissolved silica are concentrated in the hydrothermal waters compared with seawater (Table 6.6). Calcium is probably released from calcium feldspars (anorthite) as they are converted to albite by Na+ uptake, a process called albitization. Silica can be leached from any decomposing silicate in the basalt, including the glassy matrix of the rock. Globally, basalt-seawater interaction seems to provide an additional 35% to the river flux of Ca2+ and silica to the oceans.

Hydrothermal reactions involving sulphur

Seawater sulphate is removed from hydrothermal fluids, mainly by the precipitation of anhydrite (CaSO4) as the downward-percolating seawater is heated to temperatures around 150-200°C.

heat

(anhydrite) (at temperatures >150°C the equilibrium for this reaction lies well to the right)

This reaction probably consumes all of the seawater-derived Ca2+ in the fluid, and about 70% of the SO4-; if more calcium is added to the fluid from albitization reactions (see above), even more SO4- is consumed. Anhydrite formation limits the amount of SO4- entering the higher-temperature (>250°C) parts of the hydrothermal system. At these high temperatures the sulphate is reduced by reaction with FeS compounds in the basalt, and by oxidation of Fe2+ compounds, forming hydrogen sulphide (H2S) or hydrogen bisulphide (HS-). Anhydrite also forms as part of the black smoker chimney system when the hot vent fluids exit at the seabed. The hot fluids heat the surrounding seawater, which causes anhydrite to form as predicted by equation 6.18.

Most of the CaSO4 formed in the crust (and around vents) probably redissolves in the ocean bottom waters as the crust ages and cools; it thus has little effect on the overall SO4- budget of the oceans. It is well known that H2S precipitates as iron sulphide in venting hydrothermal fluids, giving rise to extensive zones of sulphide mineralization and to the 'black smoke' (Plate 6.1, facing p. 138). However, the total removal of SO42- from seawater by this mechanism is again likely to be small, since evaporite and sedimentary sulphide formation adequately removes the river flux of SO42- on geological timescales. This suggests that over long timescales much of the hydrothermal sulphide is oxidized on, or just below, the seabed.

6.4.8 The potassium problem: balancing the seawater major ion budget

The major ion budget for seawater (Table 6.2) is quite well balanced (i.e. inputs equal outputs) for all elements except K+. Laboratory studies predict that K+ behaviour will change with temperature in hydrothermal fluids. Above 150°C, in the hotter part of hydrothermal systems, K+ should be leached from basalt (Table 6.6), representing an input to the seawater budget. However, in cooler parts (<70°C) of hydrothermal systems, K+ adsorption on to altered basalt may be important, resulting in the formation of clay-like minerals such as celadonite (illitic) and phillipsite (a zeolite mineral). As there is no well-documented major removal process for K+ from seawater, it is generally believed that ridge flank low-temperature hydrothermal activity removes all of the high-temperature hydrothermal K+ input to seawater and probably some of the river flux also.

A process that might affect the K+ budget in a small way is K+ fixation during ion-exchange reactions on clay minerals. Laboratory experiments have shown that degraded micas and illites (see Section 4.5.2), stripped of their K+ during weathering, but which retain much of their layer charge, are able to fix, irreversibly, K+ from seawater. The process involves the replacement of hydrated cations for dehydrated K+ in the interlayer site, fixing the K+ in its 'mica' site (see Section 4.5.2). Globally, this process might remove another 10-20% of the K+ river flux to the oceans (Table 6.3).

The imbalance in the K+ budget and small imbalances in other budgets may be nullified by a number of processes. One possibility is the concept of 'reverse weathering reactions'. In reverse weathering, highly degraded clay minerals react with cations, HCO3- and silica in seawater to form complex clay mineral-like silicates. An example reaction addressing the K+ problem would be:

Concentration (mmol l 1) 0 20 40 60 80

0 0

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