Chemical exchange during hydrothermal alteration of basalt by seawater—I. Experimental results for major and minor components of seawater

Cctiimia Q Rrpmon

et Cormochimia Acta, Vol. 42 pp. I103 to I I IS Rcls Ltd. 1978. Pnnted in Great Britain

Chemical exchange during bydrotbermal alteration of basalt by seawater-I. Experimental results for major and minor components of seawater MICHAEL J. Morr~* and HEWRICHD. HOLLAND Department of Geological Sciences, Harvard University, Cambridge. MA 02138, U.S.A. (Received 9 November 1977; accepted in revised form 22 March

1978)

Abstract-Fresh mid-ocean ridge basalt of varying crystallinity has been powdered and reacted with seawater and an artificial Na-K-G-Cl solution at 2@I-500°C and 500-loo0 bar in sealed gold capsules. Water/rock mass ratios of l-3 were used and durations ranged from 2 to 20 months. These time periods were sufficient for most elements to approach a steady-state concentration in solution which was determined by equilibrium with alteration minerals (Mg, 50,. SO.), by rate of formation of these minerals (Na. Ca), or by depletion from the rock (K, 8, Ba). The resulting solutions closely resemble the brines from the basalt-seawater geothermal system at Reykjanes, Iceland. Mg was almost completely removed from seawater into the alteration products smectite, tremolittactinolite, or talc. Sulfate also was removed to low concentrations, both by precipitation of anyhydrite and by reduction to sulfide. Net transfer of Na from seawater into solids occurred in most experiments by formation of sodic feldspar and possibly analcime. Sr was removed from seawater in some experiments but showed no change or a small gain in others. SiO*, Ca, K, Ba, B and CO* were leached from basalt and enriched in solution. SiOt concentrations were controlled by saturation with quartz at 300°C and above. The principal Ca-bearing phases which formed were anhydrite, the hydrated Ca-silicate truscottite, tremolite-actinolite. and possibly wairakite. No K-rich phases formed. For some minerals the crystallinity of the starting basalt affected the amount which formed. Removal of Mg from seawater into solid alteration products occurred rapidly and was balanced largely by leaching of Ca from basalt. Net transfer of Na from seawater into solids occurred more slowly and was balanced mainly by leaching of additional Ca from basalt. Thus, reaction between seawater and basalt at low water/rock ratios can be considered to consist of two exchanges: Mg for Ca, and Na for Ca.

lNTRODUCTlON THE EXB~NCE of submarine hydrothermal systems along actively spreading mid-ocean ridges was first suggested by ELDER (1965) on the basis of the high

conductive heat flow measured there and by analogy with subaerial geothermal areas. Since then a large body of evidence has accumulated which indicates

that convection of heated seawater through the oceanic crust, with associated submarine hot spring activity, is very probably a common and perhaps a pervasive phenomenon along sea floor spreading anters of the world’s oceans. This evidence has been summar-&i recently by several authors (SPOONER and FYFT,1973; L~IER, 1974; WILLIAMS et al, 1974). The potential geochemical significance of hydrothermal activity along mid-ocean ridges was fist pointed out by DEFFEYES (1970). He related the occurrence of metamorphosed basaits in the oceanic crust to hydrothermal circulation, suggesting that the basalts had been altered by seawater at temperatures up to a few hundred degrees Celsius and that the process resulted in exchange of Mg, Na, K and Ca between the oceans and the oceanic crust. He further

l Present addras: Department of Chemistry. Woods Hole Oceanographic Institution, Woods Hole, MA 02543, U.S.A.

stated that the fluxes of these elements during hydrothermal alteration could be appreciable relative to the amount delivered to the sea by rivers. Deffeyes’ hypothesis has remained difficult to evaluate because neither the bulk composition nor the thickness and extent of metamorphosed basalt in the oceanic crust are well known. A second geochemical process that has been linked to submarine hydrothermal activity is the formation of Fe-Mn-rich, Al-poor sediments and crusts along actively spreading ridges (see BCETROM. 1973, and BONAITI, 1975, for refs.). Recent studies have concluded that these sediments and crusts probably were deposited from seawater which had reacted with basalt at elevated temperatures, either surficially during slow cooling of submarine lava flows (CORLISS, 1971) or at deeper levels within the oceanic crust during low-grade metamorphism (PIPER,1973; DYMOND et al., 1973). A likely analogue to mid-ocean ridge hydrothermal systems can be found on the Reykjanes Peninsula of southwestern Iceland, the iandward extension of the Reykjanes Ridge. In the Reykjanes geothermal area seawater is altering basalt to zeolite and greenschist facies at temperatures of 200-3oo”C ~OMASWN and KRBTMANN~~WITIR,1972). Hydrothermal

fluids sam-

pled from deep wells have the same chloride conantration as seawater but are enriched in K, Ca and

M. J. MOTTL and H. D. HOLLAND

1104 SiO,

and depleted

seawater

in Mg, SO0 and Na relative

to

(B&NSSON et ai, 1972).

The present study has been undertaken in order to better assess the geochemical significance of basaltseawater interaction during hydrothermal alteration of the oceanic crust along mid-ocean ridges. The specific objectives have been to determine the direc-

tion and extent of chemical exchange between basalt and seawater

under

might be encountered

a range

of conditions

within a mid-ocean

such as ridge, and

to identify those factors which could contiol the exchange in the natural setting. This information will then be used to evaluate the potential of the process for the transfer of various elements to and from seawater and for the formation of submarine hydrothermal deposits. Two other laboratories have recently performed basalt-seawater experiments with much the same objectives. BISCHOFF and DICKSON (1975) reported on a six-month experiment at 200°C 500 bar, using the Dickson hydrothermal apparatus. which allows fluid to be sampled during an experiment. Our results at this temperature agree well with theirs, as will be seen. Further work from this laboratory was reported by SEYFRIEDet al. (1975), SEYFRIEDand BI%X~FF (1977), BISCHOFF and SEYFRIED(1977) and SEYFRIED and MofTL (1977). HALGH f1975) reacted

basaltic

glass

with seawater in apparatus similar to that used in the present study and in the same temperature range. 200-500°C. His experiments were of two to four weeks duratioh and were too short for the solutions to reach a steady-state composition, but his results

agree qualitatively

with ours.

The results of our experiments will be presented in three parts. Part I describes the silicate alteration mineralogy and discusses controls on the concentrations of Na. K, Ca, M& SO, SiOz, CO*, B, Sr

and Ba in solution. ,Part II will describe the oxide-sulfide mineralogy and present data for reduced and oxidized sulfur species, pH, Fe, Mn, Cu and Zn. The implications of our results for geochemical mass balances and the origin of submarine hydrothermal deposits will be discussed in Part III. EXPERIMENTAL

DESIGN

These are closed-system experiments, yet the process about which they are meant to inform us is an open-system one. Use of a closed-system approach is justified because of the relative rates of fluid flow vs reaction between fluid and rock in many hydrothermal systems. If the reaction rate is fast relative to the flow rate through a hydrothermal system, so that unreacted fluid passing through some volume of rock is largely reacted upon leaving that volume, then the erperimental design employed here is applicable to the natural system. ‘Largely reacted’ means more or less equilibrated with the alteration mineral assemblage forming at the temperature and oressure conditions within the volume of roik. These ex~er~en~, and especially those of HAJASH119751and BISCHOFFand DICKSON119751. demonstrate that‘reaiion between basalt and seawiter lrs’ rapid at 200-XNYC, at least so long as seawater has access to fresh rock surfaces. In the experiments this access is provided by grinding the rock to a powder prior to reac-

tion. Within a natural hydrothermal system local areas will exist where access is not availabie;. for example, within a major fracture, the walls of which have been completely altered. Such localities, however, apparently are not typical of many hydrothermal systems as a whole. One of these is the Rekyjanes geothermal system. Fluids sampled from hot springs and drillholes at various depths there have

similar compositions; small differences are attributable to differences in temperature and corresponding alteration mineral facies and, for the springs, to boiling (BJORNSSON et al., 1972). Deep waters sampled from the Salton Sea geothermal system show the same uniformity (HELGESON, 1967). This ihdicates that, at Reykjanes, reaction rates at 200-300°C on the whole are rapid relative to flow through the system. The network of flow channels, the amount of rock accessible to alteration, and the residence time of the fluid within the system are such that incoming seawater largely equilibrates with the alteration mineral assemblages forming there before it leaves the system. WaterJrock

ratio

Seawater/rock mass ratios used in this study are 1, 2 and 3. These values are within the range of water/rock ratios suggested for present and past subaerial geothermal systems. The water/rock ratio for a geothermal system is defined here as the mass of water which has passed through the system during its lifetime divided by the mass of hydrotherm~ly altered rock within the system. From oxygen isotopic evidence CLAYTONand STEINER(1975) calculated a minimum water/rock mass ratio of 4.3 for the Wairakei system in New Zealand, while ESLINGERand SAVIN(1973) suggested a ratio of 0.3 for the nearby OhakiBroadlands area, and CLAYTONer al. (19681 calculated a value of 0.45 for the Salton Sea geothermal s&em in California. Taylor and co-workers have interpreted oxygen isotopic data from various shallow intrusives in terms of fossil hydrothermal systems with water/rock ratios I l-2 (see T AYLOR.1974, for refs.). The seawater/rock ratio in the Reykjanes geothermal system is less than 2 (MOITL et al., 1975; HARDIE,1976). Cr,,sra,,initl

_ Relict textures in oceanic metabasalts indicate that the crystallinity of the original rock was often an important factor in determining what mineralogical and chemical changes occurred (CANN, 1969; MELSONand VAN ANDEL. 1966; AUMENTOet al., 1971). During greenschist metamorphism glassy parts of subm~ine flows commonly are altered to chlorite. More crystalline basalt may be altered to a typical greenschist assemblage of albite, actinolite, epidote, chlorite. and sphene. Plagioclase often is albitized and olivine altered to chlorite. In these experiments the crystallinity of the starting material has been varied from g&s to hblocrystalline in order to assess the effects of this variable on the mineraloev __ and chemistrv of the product phases. PROCEDURES Experimental

Experiments were conducted in standard large-volume hydrothermal apparatus. Each experimental charge consisted of weighed amounts of rock powder and seawater or Na-K-Ca-Cl solution, with argon as an inert &lIer gas, inside a gold capsule which was welded shut. The capsules were 10 cm long and had a total volume of about 15 cm3. Six of these capsules fit vertically into one 250ml capacity pressure vessel: Each experimental run was terminated by cooling the vessel in a blast of air for IS-30 min. The vessel was then opened and the capsules were removed, cleaned and weighed. Each capsule was placed under an argon atmosphere and punctured near the top with a hypodermic needle and syringe. One milliliter of gas was

Hydrothermal

II05

alteration of basalt by seawater

removed for H2S (P4, runs 1, 2, 3) or CO2 analysis (runs 4. 6). The solution was then withdrawn into a second syringe. from which aliquots were ejected through a 0.45 pm syringe-mounted MF-Millipore” filter for analysis. Once the pressure vessel was open 20-30 min were generally required to process six capsules in this manner. The entire procedure. from the beginning of cooling to the final chemical stabilization of all solution aliquots. took from 1 to 3f hr for the six-capsule experimental runs. the chief bottleneck being the opening of the pressure vessel. The capsules were then cut open and the rock removed. Analytical Upon sampling the solutions, aliquots were taken for total sulfur and reduced and oxidized sulfur species. A third aliquot was ejected into a Beckman ‘one-drop’ pH electrode. A fourth aliquot was injected into a serumcapped syringe containing I ml of I& gas and a drop of HCI. After the aas had eauilibrated with the acidified sample solution, the solution was ejected into a volumetric flask for dilution and measurement of major elements. while the gas was analyzed for CO1 in a gas chromatograph. A fifth aliquot was diluted x 20 to stabilize dissolved silica. A sixth and final aliquot was acidified with HCI or HNO, and used for minor element analysis. Solutions were analyzed for Na, K, Ca Mg and Sr by flame atomic absorption spectrophotometry; Ba and B by DC-arc emission spectrography of salt residues left after evaporation, and comparison with similarly prepared startdards; sulfate by the method of RAYMWASHAY and HOLLAND (1968); and silica both by spectrophotometry and by AAS using a nitrous oxide-acetylene flame. Uncertainties are *2% for Na and Ca; 3% for K; 8% for Mg at -z IOppm and 2% at > IOppm; 4oj, for Sr; 20% for Ba and B; and 3% for silica. Although a single-phase fluid was present in all capsules under experimental conditions, a separate gas phase exsolved on cooling. Gas from runs 4 and 6 was analyzed for CO* by gas chromatography. The CO2 content of exsolved gas from runs 1, 2 and 3 was approximated by assuming that the gas had equilibrated with the solution at 25” and calculating the gas composition from the measured solution composition. Total carbonate in the single phase fluid under experimental conditions was then calculated by adding the solution and gas compositions together in the proper proportions. Values for total carbonate calculated in this way are judged to be accurate within +2oPi,, except for the 500°C runs. where the uncertainty is +Wh. Starting basulf The basalt used in the experiments is typical low-K oceanic tholeiite with a composition nearly identical to that given by SEWWED and BLWHOFF (1977). Its crystallinity has been varied in four steps: holocrystalline (hx), crystalline (xt), crystalline plus glass (x + g), and glassy (gl). The basalt used in most of the experiments was dredged from the median valley of the Juan de Fuca Ridge as platy fragments about 4cm thick. These plates have chilled glassy margins on two opposite sides and appear to be very fresh. The glassy margins are 2-7mm thick, lustrous black, and contain more than 905; glass. These margins were sawed off. ground, and used as ‘glassy’ starting material. The interiors of the plates were used as ‘crystalline’ starting material and consist of plagioclase microlita averaging 250 x 20fim in size, smaller pyroxene and olivine microlites, a few 1-3 mm plagioclase phcnocrysts. and about 207< glass. Qystalline plus glass’ is a 3: 1 wt mixture of the crystalline and glassy powders and is about 40% glass. The holocrystalline basalt was dredged From the Bianco Trough, near the intersection of the Blanco Fracture Zone with the Juan de Fuca Ridge. It consists of plagioclase. pyroxene, quart& opaques and a few percent

deuteric alteration products. Texturally it is subophitic to intergranular with a mean grain size of IO(t3OO~m. A Mid-Atlantic Ridge basalt from the median valley at 26”N was used in a single experiment, 3E. This sample consisted of the rim from a pillow basalt. Its crystallinity step is *x + g’. All samples were ground to a grain size of < 149 pm. RESULTS Mineralogy Description. The principal mineralogical changes which occurred during the experiments are summarized in Table I. At WC and above the rock powders became more or less lithified. taking on the cylindrical shape of the cap sule interior. The powders were originally gray in color and remain so after 20 months at 200”. but at 300” they turned very greenish. Above 300” they grade back to gray. Distinct color zones parallel to the rock-water interface can be seen in many of the lithified cylinders from the 300 and 400” experiments, a green zone near the interface grading downward into more grayish zones toward the base. Reaction with seawater for 6-20 months has had remarkably little effect on the original basalt minerals. For the holocrystalline and crystalline starting materials. the most striking feature of the whole-rock X-ray diffractograms is the similarity of the patterns recorded before and after reaction. No change is discernible in the positions or intensities of the feldspar or pyroxene peaks at any temperature. Olivine persists with little or no change in its X-ray pattern after 20 months at MO”, but has completely disappeared in all of the higher temperature experiments. For the glassy starting material, which initially showed only a broad hump in the 4.2-2.4 A region and a few ooorly defined peaks, reaction at WSOO’Y resulted in formation of a plagioclase feldspar, the X-ray diffraction pattern of which could not be distinguished from that of plagioclase in the unreacted crystalline starting material. This newly-formed plagioclase resides almost entirely in the greater than 2pm size fraction. Its composition has not been determined but it probably is albite in the 300” experiments. Plagioclase which formed in the 500” experiments may be as calcic as andesine (LIOU et al., 1974; ORVILLE, 1972). HAJASH and TIEH (1976) determined that the plagioclase which formed in their basalt glass-seawater experiments at 400 and WC was oligoclase of about 15 mole per cent anorthite. Also occurring exclusively in the greater than 2/un size fraction, in one 300” experiment only (P4). is a zeolitc identified by its X-ray pattern as either analcime or disordered wairakite (Lnou. 1970). This phase apparently required a long time to crystallize, as it is absent in similar experiments which ran for six and eight months, as compared to 19 months for P4. Moreover, the mineral formed only by reaction of basalt glass. It is absent in experiment P6, in which crystalline basalt was reacted with seawater for 17 months. The less than 2 pm fraction from the glassy basalt experiments consists mainly of smectite. Also present are tremolite-actinolite in the 400 and 500” experiments and truscottite in the 300” experiments. Truscottite. a hydrated calcium silicate with a formula near CalSi,Os.(OH)l (CHAU(ERS et al., 1964). occurs in the less than 2/rm fraction only. Smectite grew in sufficient quantities in most of the experiments to be detectable easily in X-ray diffractograms of the whole-rock powders. Above 300” the crystallinity

of the basalt noticeably influenced the amount that formed, especially in the holocrystallioe basalt experiments, where smectite peaks were very weak at 400 and SW. The smectita from all temperatures responded identically to treatment, expanding from I5 to 17A on glycolation and then collapsing to about 10 A on heating to 400°C. Their 060

M. J. MOTTLand H. D. HOLLAND

1106

Table I. Alteration minerals identified from basalt-seawater 200°C 300°C

4ooC

500°C

experiments

Smectite, anhydrite Oiivine persists for 20 months Smectite. anhydrite. aibite, anaicime of wairakite, sphene, quartz. truscottite [Ca~Si~O~(OH)~] Pyrite (except in P6 and hx ex~rjments). pyrrhotite (in P6 only). hematite?) (in hx experiments only), magnetite(?), chaicopyr~t~?) Oiivine disappears Smectite, anhydrite, aibite or oiigociase. sphene, tremoiite-actinoiite. quartz Pyrite (except in hx experiments). hematite, magnetite. chaicopyrite(?) Oiivine disappears Smectite, very rare anhydrite, oiigociase or andesine, sphene, tremoiite-actinoiite, talc, quartz(?) Hematite, pyrrhotite. magnetite(?) Oiivine disappears

hx = hoiocrystaiiine basalt as starting material.

spacing is near 1.54 indicating they are dominantly trioctahedrai. Smectites formed in the experiments using Mg-free Na-K-Ca-Cl solution have X-ray patterns similar to those from the experiments using seawater. The crystaiiinity of the basalt also influenced the formation of tremoiite-actinoiite and talc. Tremoiite-actinoiite formed much more abundantly at 500 than at 400” and, although present in experiments at ail crystaiiinities, it is much more plentiful as an alteration product of glass. At 500°C it is the dominant alteration phase in the glassy experiments. Talc, which formed only at 500”. could not be detected by X-ray in the glassy whole-rock powders. In the crystalline and hoiocrystaiiine powders, however, it is generally more abundant than tremolit~actinoiite. Quartz grew at 300 and 400” from crystaiiine basalt which originaily contained oiivine. Quartz was not detected in the experiments which used basalt glass, however. In the crystalline and glassy basalt experiments at 500’ identification of quartz is questionable. The hoiocrystaihne basalt contained quartz initially and this persisted at ail temperatures. Anhydrite was found in the whole-rock powders at ail temperatures except 500”, where its absence is attributable to nearly quantitative reduction of sulfate to sulfide. It appears to be more abundant at 200 and 300 than at 400”. Besides the formation of alteration minerals in the whole-rock powders many phases grew on the capsule wails, either on the upper portion directly from solution or on the lower third between the rock and the gold wails. Some minerals such as smectite and anhydrite grew preferentially at the rock-water interface. At 200°C the capsule walls stayed fairly clean. At 300 and 400” they were uniformly coated with about 10 and 20mg of smectite, respectively. Anhydrite crystals up to 0.8 mm long occur on top of and within the smectite layer at both temperatures, although much more abundantly at 300’. At 400” the anhydrite blades are highly corroded, leaving casts in the smectite which reveal the cIean gold wails behind. This indicates that some anhydrite crystaiiized before smectite began to form and that partiat dissolution of the anhydrite subsequently occurred. probably on cooling when the run was terminated. At SOWCthe upper capsule walls were unifo~Iy coated with 1%20mg of nearly pure tafc. Thii occurred for all crystaiiinities, including glass, and for those ex~riments which used Mg-free Na-K-Ca-Ci solution as well as for those which used seawater, although to a lesser extent. Minor amounts of tremoiite-actinoiite and smectite, and very minor anhydrite, also were detected in some cases.

Comparison

with

natural

assemblages.

The typical

in metabasites dredged from the ocean floor consists of chlorite + albite + actinolite &-epidote. Also common in these rocks are smectite, quartz, pyrite, sphene, and relicts of igneous plagioclase and pyroxene. Less common are pumpelIyite (MELSONand VAN ANDEL, 1966), prehnite, talc, ad&aria and calcite. Zeoiite facies minerals from abyssal metabasalts include analcime, wairakite, natrolite, thomsonite, chabazite, laumontite, stilbite, heulandite, smectite and mixed-layer chlor~t~m~tite (MIYA~RO et al., 1971; AUMENT~et al., 1971; JEHL greenschist

assemblage

et al., 1976).

Alteration of basalt in the Reykjanes geothermal system was described by TOMA%ZON and KRISTMANNSWI-TIR (1972) and KRISTMANNSDOTTIR (1976). The transition from zeolite to greenschist facies occurs there between 200 and 300°C. Below 200” smectite with X-ray characteristics identical with those formed in the experiments (their types 1 and 2) is the dominant sheet silicate. Chlorite first appears between 230 and 280” and replaces smectite at higher temperatures. Mixed-layer chlorite-smectite dominates at 20@-230”.The zeolites mordenite, stilbite and mesolite occur only below 230”, while analcime and minor wairakite occur at higher temperatures as well. Prehnite is present from just below the zeolite zone and toyard higher temperatures. Epidote forms as low as 200” and becomes a major phase above 260-270”. Present in varying amounts throughout the system are albite, K-feldspar, quartz, calcite, anhydrite and pyrite. Hematite occurs in the upper 150m. Of the basalt mine&s olivine is most altered and pfagiociase least, with pyroxene intermediate. Some marked differences in mineralogy are apparent between Reykjanes and the ocean-floor metabasafts, as well as between them and the experimentai products. Most notable are the presence of anhydrite and abundant K-feldspar and calcite at Reykjanes and the absence of actinohte and talc. SPOWER and

Hydrothermal alteration of basalt by seawater FYFE(1973) suggested that actinolite forms in the oceanic crust at higher temperatures than have been reached by drilling at Reykjanes (-280°C). This suggestion is supported by the occurrence of actinolite in the experiments at 400 and 500” but not at uw)“, and by the sole known occurrence of amphibole in an Icelandic geothermal area, Krafla, only at 280°C and above (KRWMANNSWYTIR. 1975). Talc formed only in the 500” experiments and its absence at Reykjanes may also be due to relatively low alteration temperatures there, or to the low seawater/rock ratio. Calcite and K-feldspar precipitate because of the high concentrations of CO2 and K+, respectively, in the Reykjanes brines. The general pattern of alteration in the experiments, with crystallinity exercising some control over the relative amounts of various alteration minerals which form, is similar to that in the natural settings where seawater interacts with basalt. Some glaring discrepancies exist in the alteration mineralogy, however. Two of the four principal minerals of the green-, schist assemblage could not be identified in the experimental products: chlorite and epidote. Absence of the latter is not surprising, since epidote has proved notoriously difficult to nucleate in hydrothermal experiments at less then 2 or 3 kbar pressure (LIOU, 1973). The absence of chlorite, along with the metastable formation and persistence of smectite even at 500°C for periods up to 9 months, is obviously a problem of kinetics. A similar metastable persistence of smectite was noted in chlorite-bearing oceanic greenstones by MIYA~HIROet al. (1971) Other discrepa;icies between t& experimental and natural mineral assemblages can be attributed to kinetic factors, such as the formation of truscottite rather than a Ca-AI silicate in the 3&Y experiments. There are, however, a number of similarities. Among the minerals that are found both in the natural settings and in the experimental products are quartz, smectite, analcime or wairakite. albite, sphene, pyrite, tremolite-actinolite and talc. Anhydrite formed in the experiments as at Reykjanes. We shall see in the next section that for the solution chemistry, the similarities are much more important than the differences. Solution chemistry Major elements. Data in Table 2 show that changes in the concentrations of major elements are fairly consistent over the entire 200-500” temperature range, K and Ca increase to several times their seawater values, While Mg and SO4 drop to low levels Removal of Mg into solid phases is particularly effective: in nearly all cases less than 37/, of the original seawater Mg remains in solution. Sulfate concentrations, although very low compared with those of seawater, are generally higher than the 20-30ppm in the Reykjanes waters. This is due to dissolution of anhydrite during the run termination, and perhaps also to oxidation of some reduced sulfur during sampling. The addition of Ca

C.C.A. 42:B-c

1107

from this source is a small fraction of the total increase in dissolved Ca. Net transfer of Na between rock and seawater is obscured by an increase in salinity which resulted from addition of several percent water to the rock. To illustrate the Na transfer we have calculated the final Na concentration relative to the initial mass of water. The difference between this value and the initial value for Na in seawater has been plotted against temperature in Fig. I, along with similarly adjusted data for three Reykjanes drillhole waters (BJORNSSON et al., 1972). Above 200°C nearly all the solutions have lost Na to the altered rock. The average loss is on the order of l@lOppm, or 10% of the original Na in the seawater. Comparison of the data in BJORNSSONet al. (1972) and MOTTL et al. (1975) with those in Table 2 shows that the solutions from the uw)” experiments are remarkably similar in composition to the deep reservoir water in the Reykjanes geothermal system. This is somewhat surprising, considering the differences in alteration mineral assemblage and the relatively brief duration of the experiments. The close resemblance of the solutions from the experiments to those at Reykjanes is important for two reasons. First, it sug gests that the experiments have closely approached a steady state with respect to fluxes of certain major elements; that is, the rates at which Na, Ca, Mg, and other elements are being dissolved from the primary phases (glass or crystalline) in basalt are nearly balanced by the rates at which these elements are removed from solution into various alteration phases. Thus, the formation of the alteration assemblage represents a significant control on the solution composition through buffering. Secondly, it implies that the rather large difference in alteration mineral assemblage between the experiments and Reykjanes has an apparentiy insignificant effect on the solution composition. The similarity of the solutions, in fact, may actually contribute to the metastable persistence of the experimental assemblage, by minimizing the driving force for reactions that would produce a stable assemblage. Thus, the problem of reaction kinetics in the experiments may not be a serious one where the solution compositions are concerned. Further evidence for the approach to steady state comes from B~XH~FF and DICKSON’S(1975) basaltseawater experiment at 200°C. Their Ca and Mg data are compared with those from Reykjanes and the present study in Fig. 2. Each of their points represents a sample withdrawn during the course of a six month experiment. Successive samples show a continuous decrease in Mg, while Ca initially drops, then rises, and finally levels off after about three months. A similar diagram for K shows that the concentration of this element reaches a constant value in only one month. Their Ca and K data agree reasonably well with that from experiments Pl and P2. which ran for 20 months at 200”. Higher Mg concentrations from Pl and P2 may be due to slower reaction or

1108

M.J. MOTTL and H. D. HOLLAND Table 2. Composition

of solutions from basalt-seawater Pl : 60.2day

200°C. 500 bar

experiments. Concentrations

P2: 582 day

No.

xtal

w/r

pH

Na

K

Ca

Mg

Sit&

3

Sr

Ba

PI P2

xt xt

I 1

4.6

I1235 II312

628 698

917 1223

200 97

633 554

4.1 7.4

3.5 5.2

0.9 0.9

5.1

300°C. 600 bar 4C P6 IA 4A P4 IC ID IE IF

2A 2F 2B 2C 2D 2E

I ! 1 1 1 I 1

hx XI xt xt gl gl x+g xl-g x-t-g

2 3

WC.

700 bar

xt xt gf x+-g x+g x+8

I

P4: 576 day

5.8

I1695

5.4

ff308 11538 11785 10659 9814 10765 10894 11084

58 5.75 6.0 5.6 5.6 5.35

3E 3D

hx xt xt gl x+g x+g x+g

1

4.1 3.9

1

3.9

1 2 3

3.6 3.8 3.85

Starting seawaters

1 I

P6: 506 day

I185 2.5 2003 I.5 1626 10.9 1374 3.8 3073 38 3263 2.5 2243 6.9 1983 3.7 1856 5.6

Run No. 1: 236 day

851 61.5 400 lOI5 494 39a 395 415 416

5.9 7.7 5.9 4.6

5.2

2.9 8.5 7.7 6.7 5.9 4.2 7.8 5.8 5s

1510 1650 1667 1678 1722 1684

6.8 7.7 8.0 7.1 6.8 6.5

8.2 7.8 6.3 8.9 7.9 6.9

2 I.8 2.8 5

so:-

ecu:

290 340

zco:*

r41 >43

Run No. 4: 172 day 20f 270 197 169

0.7 I.5 1.6 2.0

357 298 272 360 208 XI 204 60 134

3.6 3.1 2.0 2.2 2.0 2.1

477 260 544 531 595 543

500 680 603 615 654 573

131 t@G 52 107 83

912 519 840 970 555

614 630

708 627

223

196

I48 207 165 159

Run No. 2: 272day

5oo“C, 1000 bar 3F 3A 6A ::

9cll 1346 1240 1158 1301 1026 i255 886 726

in ppm

11829 11601

1113 1181 108f0 11% 11813 1144 11113 895 11156 762

1249 12.4 Ii67 9.5 2084 10.7 1526 7.8 1551 7.9 1565 10.5

Run No. 3: 268day,

1 1I

4,o 3.3 3.55 3.3 3.5

11322 9617 lO833 10328 9777

I 3

3.2 1.9

10787 11435

1107 425 1244 91 f I182 887 1137 1211 1165 1040 990 753

Run No. 6: 167 day

14.6 23.5 39.5 33.1 29.2

1832 2258 2531 2000 2257

9.9 10.8 5.5 9.0 8.0

0.2 1.2 2.2 2.5 3,4

3.8 2.0 -7.2 7.3 3.3

1024 26.7 1106 16.7

2080 2181

7.7 7.0

3.5 4.I

1.9 1.9

Atlantic: ‘P’ runs

Sargasso: alt others

XC&

600

Cl

Atlantic 10490 389 4% 1286 4.1 6.8 2650 I20 19010 Satgasso 11070 403 424 1340
to back reaction on cooling. but the low values found both at Reykjanes and in BischofT and Dickson’s experiment again indicate approach to steady state in a relatively short time. Also &own in Fig. 2 are the Ca and Mg data from the 300’ basalt-seawater experiments. The very low Mg concentrations relative to seawater, and the resemblance of the experimental solutions to the Reykjanes waters, are obvious. It the experimental solutions are near steady state and their compositions are largely controlled by formation of the alteration assemblages. it should be possible to approach the same composition from the opposite direction. To test this supposition a number of experiments were performed using a Na-KXa-CI solution instead of seawater. Although the total mola-

lity was similar. initial concentrations of K and Ca in this Mg- and SO,-free solution were greater than those attained in the basalt-seawater experiments run earlier. The results, given in Table 3. show that the directions of net flux for Na. Ca, Mg and SO4 and. in some cases. K have reversed: Ca is now depleted in solution while Na. Mg and SO., are enriched. The reversal of the Ca flux is illustrated in Fig. 3. The close approach to a steady-state Ca concentration. bracketed from opposite directions in all cases except one. is evident. although the actual concentration varies at each temperature with the crystallinity of the basalt. The Mg and SO, concentrations in the reacted solutions likewise fall in the same range whether the initial solution was seawater or the Mgand Sod-free water. The K data are more ambiguous

Hydrothermal

alteration of basalt by seawater

1109

ing 400”. The seawater/rock ratio has an effect on the Na and Ca concentrations also, but only at 300”. and there its influence is slight. An analogous diagram for the change in K concen-

see

400

300 T,

1

OC

Fig. 1. Change in the concentration of sodium in solution resulting from transfer of sodium between basalt and seawater during hydrothermal alteration at Reykjanes (triangles) and in the experiments (circles and squares). Open square = holocrystalline. x-square = crystalline, xcircle = crystalline + glassy, open circle = glassy basalt. Lines connect points for experiments with similar crystallinity and duration.

tration is given in Fig. 5. K, like Ca. is removed from basalt into seawater during interaction at 20&5OO”C. Unlike Ca, however, it shows no clear pattern due to crystallinity of the basalt. Instead, the water/rock ratio of the experiments has exerted a strong control on the K increase in seawater reacted at 3O&soo”. The increase is inversely proportional to the water/ rock ratio. This is undoubtedly due to the very low initial K content of the basalt. The K increase shown

in Fig. 5 indicates that a significant fraction of the total K in the basalt has been removed. Of the 150Oppm in the rock initially, about half has gone into solution at 1: 1 water/rock. At 2: 1 and 3: 1 the figure is two-thirds. Curiously, the leaching has been equally effective at 300, 400 and 500” for a given water/rock ratio. suggesting that most of the readily

showing a clear reversal ments.

These

results

only in the glassy experidemonstrate that the solutions

in the basalt-seawater experiments are near steady state. The rates of removal of some elements apparently vary, however,

with crystallinity

OOA

’

of the basalt.

The effect of crystallinity on the Ca concentration is shown in more detail in Fig. 4. Here the change in Ca concentration, adjusted to the initial mass of seawater, is plotted against temperature, as was done for Na in Fig. 1. In every case except one the Ca content of the solution has increased at the expense of the basalt, generally by 1OOOppm or more. The amount of increase varies inversely with crystallinity: the more glassy the basalt, the greater the increase. At higher temperatures the effect becomes less pronounced. The mirror image of the crystallinity effect for Ca can be seen in Fig. 1 for Na, up to and includTable 3. Composition

D --

Fig. 2. Calcium vs. magnesium in seawater which has reacted with basalt at 200-300°C in the experiments and at Reykjanes. Magnesium decreased continuously with time in BISCHOFFand DICKSON’S(1975) experiment.

of solutions from basalt + Na-K-Ca-CI-Hz0

experiments. Concentrations

in ppm

3OO”C,600 bar: 172 day No.

xtal

4D 4B 4E

hx xt gl

w/r 1 :

pH

Na

K

Mg

SiOl

B

Sr

Ba

SOi-

2.2 2.9 0.9

862 1075 749

4.2 3.7 3.6

3.6 10.0 5.6

4 5 1.5

48 74 41

234 203 170

289 230 194

738 15.3 1316 45.8 1732 47.5

2149 2638 2755

2.5 3.7 5.1

1.3 3.7 5.3

8.3 7.5 4.5

t 22

186 415 642

664 382 429

0

0

0

0

0

Cl 23070

Ca

5.7 5.75 6.05

10109 10765 9842

2041 2513 2050 2509 1703 4210

3.15

11430 10580 10332

2084 2100 2011

X0$

zcor*

500°C, 1000 bar: 167day 6D 6B 6C

hx xt gl

Starting solution: Na-K-C&Cl-H,0

1 1 1

9200

2000 4000

0

‘xtal’ = ~ys~~linity of starting basalt (see text). ‘w/r’ = water/rock mass ratio. * calculated: for single-phase guid present under run conditions. Sum of measured CCOa in liquid plus calculated CO2 in gas. ** measured: for single-phase guid present under run conditions. Sum of measured ECOz in liquid plus measured COz in gas.

1110

M. J. MOTTL and H. D. HOLLAND

4eoQ’

me-

Loo-

3000.

fi

pwn

.

+r&o-

CP

-4

zone.

200-

oL

T, *C

Fig. 3. Calcium concentration in experimental solutions resulting from reaction of basalt with seawater (424ppm Ca) and Na-K-Ca-Cl solution (4oooppm Cal. Symbols are those in Fig. 1.

.

2eo

100

3oe

so0

T. "C

Fig 5. Change in the concentration of potassium in sofution resulting from transfer of potassium between basatt and seawater during hydrot~rmai alteration. Symbols are those in Fig. 1. Points joined by dashed (not dotted) lines are for experiments at 2:l and 3:l water/rock (wt); all others are for experiments at 1: 1.

runs until two days after the runs were terminated. Some silica probably precipitated during this time. Nonetheless, silica ~on~ntratio~ in Pl and P2 are still higher than for quartz saturation, and no quartz available K has been removed at each temperature was detected in the run products. It is clear that equiand that its availability is what determines the K conlibration with quartz was not reached at 200” even centration in solution. in 20 months’ time. By contrast, all geothermal waters Silica. Data for silica in Tables 2 and 3 can be circulating through basalt in Iceland at 180°C and compared with the following quartz solubilities under above have equi~brated with quartz (ARN~~N, conditions of the experiments: 1974). At 300”, quartz crystallized and grew during the 200°C 500 bar 300ppm runs in at least some experiments. Only run 4 gives 300°C 600 bar 850 ppm dissolved silica values close to quartz saturation, how400°C 700bar 1500ppm ever. Run 1 took longer to process than the others SOW’C, iCOObar 24OOppm (KENNEDY,1950). (34 hr) and it may be that some SiO, precipitated At 200”, the 8OOppm silica in BischolI and Dickbefore the samples could be diluted. son’s experiment at six months was close to satuQuartz also formed at 400 and SW, and the silica ration with amorphous silica (*9OOppm: FOURNIER concentrations in these solutions are close to equiliband ROWE, 1966). The values reported here for Pl rium values. and P2 are lower than that (550-650ppm), but this Carbon dioxide. The concentration of COz in the is probably because the dilutions which stabilized disaltered seawater of the Reykjanes geothermal system solved silica for analysis were not made for the ‘P is about 2000 ppm, BJORNSJN eral.(1972) considered two possible sources for this CO*: a juvenile sourer and leaching from the rock during alteration. Data for total CO1 in the experimental solutions (Tables 2 and 3) indicate that sizable amounts of CO2 have been removed from the rocks used in this study. The amount removed is strongly temperature dependent and a maximum concentration of nearly 1OOOppm was attained in the 500” solutions. Are the amounts removed from basalt in the experiments adequate to account for the high CO2 at Reykjanes? For the 200-308” experiments the answer is certainly no: the 200ppm in solution at 300* is an order of magnitude too low. Moreover, the total COZ a0 See concentration in the reacted solutions appears to be T. OC independent of the water/rock ratio, suggesting that Fig. 4. Change in the concentration of calcium in solution it is not controlled by availability of COZ in the rock resulting. from transfer of calcium between basalt and seaor by rate of leaching. Thus, neither a longer experiwater during hydrothermal alteration. Symbols are those in Fig. 1. ment nor one with a lower water/rock ratio is likely

-1

Hydrothermal alteration of basalt by seawater

0

most of the experiments but gained in others. The Reykjanes waters have gained more Sr than any of the experimental solutions. Thus, the direction of net Sr transfer is ambiguous, although the amounts in all cases are small relative to the 100 or more ppm found in most ocean floor basalts. The Sr concentration in solution presumably is determined by rate of leaching or incorporation into alteration products, or both.

B

lm

q 0

-1

100

1111

300

400 1.

500

oc

Fig. 6. Change in the concentration of boron in solution resulting from transfer of boron between basalt and NaK-C&Cl solution (dashed symbols) or seawater (others) during hydrothermal alteration. Symbols are those in Fig. I.

DISCUSSION: CONTROLS ON SOLUTION COMPOSITlON Element exchiges

Changes in the major element composition of seawater resulting from high-temperature interaction with basalt are summarized in Fig. 8. The changes which occurred in the experiments and at ,Reykjanes have been adjusted to their values relative to the initial mass of seawater in order to remove the effects of rock hydration and, for Reykjanes. boiling and dilution by meteoric water. The remaining change is due only to transfer of ions between rock and SOIUtion. These values have been plotted in milhequivalents/kg of solution to emphasize the relationships among major ions which result from the requirement that charge must balance in solution. In Fig. 8. points for cations which fall below the zero line represent loss of that cation from solution into the solid phases. Cations which plot above zero represent gains to solution from basalt. In order to stress the electrical neutrality requirement the diagram has been constructed such that, for a given solution, the sum of all the ions which plot above zero equals the sum of those below zero. Accordingly SO:-, the long anion whose change in concentration is important for the charge balance, has been plotted as a negative: although it has been lost from seawater it occurs above the line, because loss of an anion requires loss of a cation to balance charge. For simplicity, only the crystalline and glassy basalt experiments at 1: 1 water/rock have been included, along

to produce the 2000ppm found in the brines at Reykjanes. It follows that the CO, concentration in the solutions from the experiments may be controlled by saturation with some carbonate phase which grew during the runs but was not detected. If the experimental solutions are saturated with calcite, as the Reykjanes solutions presumably are, then the similar Ca concentration and higher CO1 at Reykjanes imply a lower pH there than in solutions from the experiments. A large input of CO2 from a juvenile or magmatic source could lower PI-I and account for the high CO1 concentration at the same time. The results of the experiments favor such a source for much of the CO2 at Reykjanes Boron, strontium and barium. Concentrations of B, Sr and Ba in solutions from the experiments are given in Tables 2 and 3, and from Reykjanes in Morn et al. (1975). As seawater contains a few ppm each of B and Sr, the net transfer of these elements between rock and solution is shown in Figs. 6 and 7. In nearly all cases a few ppm each of B and Ba have been leached from basalt into solution. Like K, both B and Ba are present in unweathered oceanic tholeiites in very low concentrations, usually less than 10 ppm. The amounts removed thus represent significant proportions of the total present. The precision of the analyses is too poor for a water/rock effect *mSr2+ to show, as in Fig. 5 for K, but it can be expected ppm that higher water/rock ratios will result in smaller increases in solution, because the increase is limited by the amount available in the rock. Ba concentrations in the Reykjanes waters are especially htgh-up to 10 ppm-presumably because of a higher Ba concentration in the basalts of the Reykjanes n I . peninsula (HART and SCHILLING,1973). For reaction 300 400 so0 with abyssal tholeiite, such high concentrations would T, OC be possible only at relatively low water/rock ratios. Fig 7. Change in the concentration of strontium in soluThe pattern for net transfer of Sr plotted against tion resulting from transfer of strontium between basalt temperature (Fig 7) parallels that for Ca (Fig 4) Unand seawater during hydrothermal alteration. Symbols are like Ca. however. Sr has been lost from seawater in those in Fig. 1.

1112

M. J. MOTTLand H. D. HOLLAND

range of ex~rimental parameters is that each is closely limited by depletion. K, as shown earlier, is nearly depleted from the rock, whereas Mg and SO4 are nearly depleted from seawater. Sulfate concentrations are held at these low levels by the extremely small solubility of anhydrite at elevated temperatures. The large loss of Mg under all conditions indicates that Mg is strongly partitioned into solid phases during high-temperature basalt-seawater interaction. At 20&5oo”C and water/rock ratios of one to ten Mg removal from seawater occurs rapidly and in every experiment has been nearly complete, often in less than two weeks’ time (HAJASH, 1975; BISCHOFF and DICKSON, 1975). Other changes in solution composition that must accompany Mg removal in order to balance charge must occur equally rapidly. Inspection of Fig. 8 shows that Mg loss is compensated at 2oo-400” mainly by loss of SO4 and gain of Ca, with a minor contribution from K and, at 200”, Na. Loss of seawater bicarbonate is negligible for the total Fig. 8. Changes in the composition of seawater which has balance, as is gain in H+, although the latter unreacted with basalt, resulting from transfer of major ions doubtedly speeds the dissolution of reacting phases. between rock and solution during hydrothermal alteration at Reykjanes and in the experiments. Lines connect points Fe and Mn become important only at 500”. Loss of for experiments of similar duration (and Reykjanes), SO,, which accounts for about half of the Mg charge denoted in months for some points at 300”. Elements and deficit, occurs largely as anhydrite. With respect to species are calcium (large circles), sulfate (inverted basait, therefore, it represents additional leaching of triangles), potassium (triangles), sodium (squares), magnesium (hexagons). and iron plus manganese (small circles). Ca. Thus, Ca is by far the dominant cation involved Solid squares and circles respresent glassy, and open in balancing Mg removal from seawater. If only squares and circles crystalline, basalt experiments. See text charged species are considered, the overall reaction for details of construction. removing Mg and forming silicates can be approximated as an exchange of Mg for Ca. The experimental evidence in Fig. 8 suggests that with three Reykjanes drillholes. Where several soluMg removal from seawater occurs far more rapidly tions at one temperature have similar values for an than Na removal. At 300”, for example, seawater ion, only one point has been plotted which represents the average. This has been done for K, Mg and SO4 reacted with crystalline basalt for six months has lost nearly all its Mg and virtually none of its Na. For throughout, and for Na and Ca in some crystalline these conditions Na removal becomes appreciable experiments. As we have seen, the general trends at 20@-500” only at 8-20 months. Where net transfer of Na from seawater into solids has occurred, its loss from soluare loss of Mg, Na and SO., from seawater and gain tion clearly is balanced by an additional gain in Ca, of Ca and K. What Fig. 8 shows in striking fashion except at 500”. This Na for Ca exchange, superimis the reciprocal relationship between Na and Ca in posed on the Mg for Ca exchange, is what gives rise this temperature range. Whereas the gain in K and the losses of Mg and SO4 are nearly constant for to the ‘mirror image’ effect in Fig. 8. Thus, basaltall the experimental solutions and for Reykjanes, re- seawater interaction at 20040” and relatively low water/rock ratios ( I 3) can be considered to consist gardless of temperature-pressure, crystallinity, and of two steps: a Mg for Ca exchange which proceeds run duration, Na and Ca show a strong dependence rapidly and nearly to depletion of seawater Mg, and on each of these variables. Because they are the only a much slower Na for Ca exchange. ions in the range 2C#400° whose change in concenOne feature in Fig. 8 that may be explained in tration varies appreciably (on the scale of Fig. 8), electerms of this two-step model is a large discrepancy trical neutrality requires that their patterns be mirror images of each other. This is the case at 200-SO@‘, between the 200” experimental solutions and waters from only slightly higher temperatures at Reykjanes. except that at 500” Fe and Mn become major cations, It seems likely that the kinetics of the necessary reaccausing a downward shift in the position of the mirror tions were simply too slow in these 200“ experiments plane. The holocrystalline basalt experiments fit the for the Na for Ca exchange to proceed. Na actually pattern of Fig. 8 except for slightly lower Fe and Mn appears to have increased slightly in the solutions. con~ntratio~ at 5&Y’,resulting in a smafl Na gain The absence of any alteration product other than rather than a loss. The reason for the nearly constant amount of smectite and anhydrite is consistent with this suggeschange in K, Mg and SO., concentrations over a tion; the formation of Na-rich phases and the adjust-

Hydrothermal alteration of basalt by seawater ment of the Ca/Na ratio in solution that should have accompanied it apparently did not occur. Silicate equilibria Demonstration that the experimental solutions are near steady state implies that mineral equilibria are important in controlling the composition of the solutions. However, the extent to which element concentrations in solution are fixed by various equilibria may vary with both the temperature and crystallinity of the experiments. Had final thermodynamic equilibrium been achieved, of course, both solution composition and mineralogy would have depended only on temperature, pressure, and the bulk composition of the system, set by the water/rock ratio. Final equilibrium was not achieved, as both unreacted, unstable phases and metastable phases are present at the end of the experiments. Thus, the steady-state concentration of an element in solution is a function of its rate of dissolution from reacting phases as well as its rate of removal into product phases, By definition the two rates are equal at steady state, but either may act as the dominant control, depending on the specific element and the experimental parameters. In the case of Mg, its removal rate is very fast under all conditions, and it is fair to conclude that its concentration is controlled by equilibrium with a Mg-bearing product phase or phases. At 200 and 300” this is smectite. In the experiments at 300” and above smectite presumably performs a function similar to that of chlorite in a natural greenschist assemblage. Smectite is joined in the 400-500” experiments by tremolite-actinolite and at 500” by talc. Mg conwntrations increase slightly over this temperature range. In contrast to Mg, Na and Ca concentrations in the experimental solutions are strongly influenced by reaction kinetics. It can be seen from Fig. 8 that the Na for Ca exchange has proceeded further in experiments using more glassy basalt. This is true for every crystallinity step at 300 and 400” and for most at XXJ’, as comparison of Figs. 1 and 4 shows, although the effect diminishes toward higher temperatures. Moreover, at 300-500” the Na for Ca exchange has proceeded further in the glass + seawater experiments than the reverse, Ca for Na exchange has gone in the holocrystalline basalt + Na-K-Ca-Cl experiments. This means that the Ca/Na ratio in the reacted solution is greater for the glassy basalt experiments than for the holocrystalline basalt experiments at the same temperature, regardless of whether the initial solution had Ca/Na lower than the final value (seewater) or higher (Na-K-CaCI solution). This effect is best illustrated in Fig. 3, which shows the zone of overlap for experiments at various crystallinities which resulted when final Ca concentration was approached from opposite directions. An analogous diagram for Ca/Na vs temperature and crystallinity would show exactly the same effect. These observations can be explained aa follows: equilibrium among various Ca- and Na-bearing sili-

1113

cates and a 0.5 molal chloride solution under the conditions of the experiments is such that the Ca/Na ratio in solution is greater than that in seawater. The Ca/Na ratio at equilibrium is also greater than that in reacted seawater which has gained Ca via the kinetically rapid Mg for Ca exchange alone. A Na for Ca exchange between solids and solution results, but the rate at which Na is taken up into product phases is much greater for basalt glass than for more crystalline basalt. At 300” Na uptake probably occurs by growth of albite, and possibly analcime as well. At 400-500” sodic plagioclase may precipitate, or cation exchange equilibrium may result in Na enrichment of pre-existing plagioclase. As net Na uptake occurs, Ca is dissolved at an equal net rate. Its concentration increases in solution until the rate of Ca uptake into solids equals its rate of solution; in other words, steady state. At this point the net rate of Na uptake equals the net rate of Ca uptake (= O), in miiliequivalents, and the actual concentrations of Na and Ca are determined by the rate of formation of Na-bearing phases. These phases grow faster from glass, and a higher Ca concentration is achieved before the net rate of Ca uptake can match that of Na uptake. The result is that the higher the glass content of the basalt, the greater is the Ca/Na ratio in the corresponding solution. Another consequence is that Na-rich product minerals such as albite and analcime grow more plentifully in the glassy basalt experiments, a fact which is borne out by their alteration mineralogy. The plagioclase exchange experiments of ORVILLE (1972) are pertinent here. Orville found that equilibrium was approached more rapidly from a high Ca/Na fluid, by way of Ca for Na exchange, than from the opposite direction which involves albitization. This implies that the Ca/Na ratios obtained from the Na-K-Ca-Cl experiments may be closer to ‘equilibrium’ values than those from the seawater experiments. The phenomenon of greater Na uptake into glassy rather than crystalline basalt is contrary to what has been described in greenstones from the ocean floor. CANN (1969) found that glassy margins of pillows from the Carlsberg Ridge had been converted completely to chlorite, while more crystalline interiors underwent Na-enrichment through albitization of plagioclase. He explained this in terms of the reiative kinetics of ditrusion and nucleation. It is apparent that the kinetics of alteration in the present experiments dialer considerably from those which prevailed for the Carl&erg Ridge rocks. This probably is the result of differing seawater/rock ratios, a subject which will be taken up further in Part III. Depletion

While the concentrations of most elements in seawater which has reacted with basalt are controlled in large part by mineral equilibria, there are a few elements whose concentrations are limited ultimately by their depletion from the rock. ELLIS(1970) referred

M. J. Morr~ and N.

it14

to these as the ‘soluble elements’. They are characteristically present in low concentrations in the rock, and are strongly partitioned into the solution rather than the solid phases. The present experiments suggest that B, K and Ba belong in this category. Boron was identified by Ellis as a soluble element in most natural thermal waters, but it is somewhat unusual for K and Ba to behave in this way. That they do so is a consequence of their unusually low concentrations in oceanic tholeiites. It can be expected that K and Ba will no longer act as soluble elements when the water/rock ratio is much lower than those used in the experiments. Under these conditions sufficiently high concentrations will be attained in solution to saturate with some K- or Ba-rich phase, such as K-feldspar or barite. This has happened for K in the Reykjanes geothermal system, where K-feldspar occurs sporadically throughout. Barite saturation apparently has not been reached there. Figure 5 shows that the K increase in seawater reacted with basalt at 200°C is much smalier than that at 300-500“. This suggests either that K is leached very slowly at 200” or else that it does not behave as a soluble element at this temperature. The dissolved K concentration in these 200°C experiments is similar to that in BISCHOFFand DICKSON’S(1975) experiment, despite their use of a water/rock ratio of ten rather than one. The agreement can be attributed to their use of BCR-1, a Columbia River flood basalt with 1.68% K20 (14,000 ppm K), rather than a low-K tholeiite, and it reinforces the notion that at the low water/rock ratios used, the K concentration in the 200” experimental solutions was controlled by mineral equilibria. Have metamorphosed basalts and gabbros from the oceanic crust been depleted in K, Ba, and B? Analyses suggest that they have for K: many greenschist and amphibolite facies metabasites have K,O contents in the range 0.06% $- 0.03 vs 0.1-0.3% for fresh basalts (B~NA~TI et al., l975; THOMPSON and MELSON,1972; CANN, 1969; MELSON et al., 1968; MELSON and VAN ANDEL. 1966). As in the experiments, most but not all of the K has been leached from the rock. Data for Ba and B are ambiguous, however, because these elements often are near the detection Iimits of standard analytical techniques and because both may be enriched during subsequent weathering on the sea floor (~~~P~~s and THObm5N, 1978). CONCLUSIONS The experiments performed in this study were only partially successful in producing the alteration mineral assemblages found in hydrothermally altered basalts and gabbros from the oceanic crust. The experimental solutions, however, are very similar to the deep reservoir fluid in the basalt-seawater geothe~~ system at Reykjanes, Iceland, where the alteration assemblages resemble those from submarine greenstones in most respects. Thus, the discrepancies

D. HOLLAND

between the experimental and natural mineral assemblages appear to be relatively insignificant for the chemistry of the solutions. The experiments were of sufficient duration for the solutions to closely approach steady-state compositions, which were controlled largely by formation of the alteration mineral assemblages. The experimental solutions are therefore a reasonably accurate representation of solutions that would be produced at similar temperatures and pressures in natural basalt-seawater hydrothermal systems with low water/rock ratios. Whether the solutions resemble the altered seawater in submarine hydrothermal systems depends upon the extent to which the experimental conditions accurately represent those which prevail in the oceanic crust. Besides temperature, the most criticai variable is undoubtedly the water/rock ratio. WOLERY and SLEEP(1976) pointed out the difficulty of using experimental data directly to calculate elemental fluxes resulting from hydrothermal alteration of the oceanic crust by seawater. They suggested that water/ rock ratios in sub-seafloor hydrothermal systems may be greater than those used in this study. This issue, while outside the scope of the present paper, is a critical one, and it will be taken up further in Part III of this series. Acknowledgements-We wish to thank R~MMUND F. CORR for performing many of the analyses, and S~EFANARN~R% SON of Orkustofnun, Reykjavik, who showed the senior author how to sample geothermal wells in Iceland. Basalt was obtained from W. G. MELSONof the Smithsonian Institution and from R. B. Sco’rr of Texas A & M University courtesy of the NOAA Tram-Atlantic Geotraverse Project. Figures were drafted gratis by WILMAWETTERSTROM. This work was part of the senior author’s Ph.D. thesis at Harvard University and was supported by Grant GA-40052 from the National Science Foundation. This paper was written in part at Stanford University with the support of NSF Grant ID 074-12880 and was revised with support of the Woods Hole Oceanographic Institution.

REFERENCES ARN~RSSON D. (1974) The composition

of thermal fluids in Iceland and geological features related to the thermal activity. In Geodynamics of Iceland and the North Atlantic Ark (editor-L. Kristjansson), pp. 307-323. Reidel. AUMENTO F.. LONCAREVIC B. b. and Ross D. I. (1971) Hudson Geotraverse: geology of the Mid-Atlantic Ridge at 45”N. Phyl. Trans. Roy. Sot. Lond., Series A 268, 623-650. BI~CHOF~J. L. and DICKSON F. W. (1975) Seawater-basalt intera~ion at 200°C and 500 bar: implications for origin of sea floor heavy-metal deposits and regulation of seawater chemistry. Earth BtsCHoFF J. L. and SEYF~~D W. E. (1977) Seawater as a geothermal fluid: chemical behaviour from 25 to 350°C. Proc. of the Second Int. Symp. on Water-Rock Interaction, I.A.G.C., Strasbourg, France, pp. IV 165-IV 172. BJORN~KIN S., ARN~RSSONS. and TOMA~ON J. (1972) Economic evaluation of Reykjanes thermal brine area, Ic&nd. Bull. Am. Assoc. Petrol. Geooiogists 56, 2380-2391. B~NATTIE. (1975) Metatlogenesis at oceanic spreading centers. In Annual Review of Earth and Planetary Sciences (editor F. A. Donath), Vol. 3, pp. 401-431. Annual Reviews.

Hydrothermal

alteration of basalt by seawater

BONATTIE., HONNOREZJ., KIRSTP. and RAIXCATIF. (1975) Metagabbros from the Mid-Atlantic Ridge at 06”N: contact-hydrothermal-dynamic metamorphism beneath the axial valley. J. Geof. 83, 61-78. Bosraosr K. (1973) The origin and fate of ~erromang~oan active ridge sediments. Acta Uniu. Stockholm 27, 149243.

CANNJ. R. (1969) Spilites from the Carl&erg Ridge, Indian Ocean. J. Petrol. IO, l-19. CHAL~RS R. A., FARMERV. C.. HARKERR. I., KELLY S. and TAYLORH. F. W. (1964) Reyerite. Minerai. Mug. 33, 821-840.

CLAYTONR. N., MUFFLERL. J. P. and WHITE D. E. (1968) Oxygen isotope study of calcite and silicates of the River Ranch no. 1 well, Salton Sea geothermal field, California. Am. J. Sci. 266, 968-979. CLAYTON R. N. and STEINERA. (1975) Oxygen isotope studies of the geothermal system at Wairakei. New Zealand. Geochim:Cosmochim.~ Acta 39, 1179-l 186. Co~~rss J. B. (1971) The oriain of metal-bearintz submarine hydrothermal solutions, 57 Geophys. Res. 76:8128-8138. DEFFNES K. S. (1970) The axial valley: a steady-state feature of the terrain. In The Megafecro~ics of Co~r~ne~ts and Oceans (editors H. Johnson and B. L. Smith), pp. 194-222. Rutgers University Press. DYMONDJ., CORLLSS J. B., HEATH G. R., FDELDC. W., DASCH E. J. and VEEH H. H. (1973) Origin of metalliferous sediments from the Pacific Ocean. Bull. Geot. Sot. Am. 84, 3355-3372.

ELDERJ. W. (1965) Physical processes in geothermal areas. In Terrestrial Heat FIow (editor W. H. K. Lee), Am. Geophys. Union Monograph No. 8, pp. 211-239. ELLIS A. J. (1970) Quantitative interpretation of chemical characteristics of hydrothermal systems. Geothermics, Spec. Issue 2, 516528.

ESLINGERE. V. and SAVIN S. M. (1973) Mineralogy and oxygen isotope geochemistry of the hydrothermally altered rocks of the Oh&-Broadlands New Zealand geothermal area. Am. J. Sci. 273, 240-267. FOURNIERR. 0. and ROWEJ. J. (1966) Estimation of underground tem~ratur~ from the silica content of water from hot springs and wet-steam weils. Am. J. Sci. 264, 685-697. HAJASHA. (1975) Hydrothermal processes along mid-ocean ridges: an experimental investigation, Contrib Mineral. Petrol. 53, 205-226.

HAJASHA. and TIEH T. T. (1976) Ex~rimentaI seawater/ basalt reactions, 200-5oo”C, 500-2000 bar: additional mineralogical data (abstract). Abstracts with Programs, Geol. Sot. Am. Ann. Meetings 8, 89Q-900.

HARDIE L. A. (1976) Origin of CaCl,-rich geothermal brines by basalt seawater interaction beneath mid-ocean ridges: some simple mass balance calculations (abstract). Abstracts with Programs, Geol. Sot. Am. Ann. Meetings 8, 903.

HART S. R. and SCHILLING J. G. (1973) The geochemistry of basalt5 from Iceland and the Reykjanes Ridge. Carnegie Inst. Wash. Year Book 72, 259-262. HELGE~~NH. C. (1967) Solution chemistry and metamorphism. In Researches in Geochemistry (editor P. H. Abelson), Vol. 2, pp. 362-404. Wiley. HUMPHREY S. E. and THOMPSONG. (1978) Trace element mobility during hydrothermal alteration of oceanic basalts: Geochim. Cbsmochtm. Acta 42, 127-136. JEHL V., PITY B. and WEBBROD A. (1976) Hydrothermal metamorph~m of the oceanic crust in the North Atlantic Ocean (abstract). EOS, Trans. Am. Geuphys. Union 57, 597-598. KENNEDYG. C. (1950) A portion of the system silica-water. Econ. Geol. 45, 629-653. KRISIXANNSDOTTIR

H. (1975) Hydrothermal alteration of basaltic rocks in Icelandic geothermal areas. Second

1115

U.N. Symposium on the Development and Utilization of Geothermal Resources, San Francisco, pp. 441-445. KR~~TMANNSDOTTIR H. (1976) Types of clay minerals in hydrotherm~ly altered basaltic rocks, Reykjanes, Iceland. Jiikull 26, 30-39. Lrou J. G. (1970) Synthesis and stability relations of wairakite, CaA12Si40i2 *2HsO. Contrtb. Mineral: Petrol. 27, 259-282.

LIOU J. G. (1973) Synthesis and stability relations of epidote. Ca,Al,FeSi,O,,(OHf. J. Petrol. 14 381-413. Lieu J: G.,-K&~&HI~S. anh ITO K. (1974) Ex~rimental studies of the phase relations between greenschist and amphibolite in a basaltic system. Am. J. Sci. 274, 613-632. LESTERC. R. B. (1974) On the penetration of water into hot rock. Geophys. J. Roy. Astron. Sot. 39, 465509. MELSONW. G. and VAN ANDELTJ. H. (1966) Metamorphism in the Mid-Atlantic Ridge, 22”N Latitude. Mar. Geol. 4, 165-186.

MELVIN W. G., THOMPSONG. and VAN ANDEL TJ. H. (1968) Volcanism and metamorphism in the MidAtlantic Ridge, 22”N Latitude. .I. Geophys. Res. 73,59255941. MIYA~H~RO A., SHOW F. and EWING M. (1971) Metamorphism in the Mid-Atlantic Ridge near 24 and 30”N. Phi!. Trans. Roy. Sot. Lond., Series A 268, 589-603. MOTTLM. J., CORRR. F. and HOLLANDH. D. (1975) Trace element content of the Reykjanes and Svartsengi thermal brines, Iceland (abstract). Abstracts with Progr~, Geof. Sot. Am. Ann. Meetings 7, 1206-1207.

ORVILLEP. M. (1972) Plagioclase cation exchange equilibria with aqueous chloride solution: results at 700°C and 2000 bar in the presence of quartz. Am. J. Sci. 272, 234-272.

PIPER D. Z. (1973) Origin of me~lliferous sediments from the East Pacific Rise. Earth Planet. Sci. Left. 19, 75-82. RAYMAHASHAY B. C. and HOLLANDH. D. (1968) Composition of aqueous solutions in equilibrium with sulfides and oxides of iron at 350°C. Science 162, 895-896. SEYFRIERW. E. JR., BISCHOFF J. L. and DICKSONF. W. (1975) Basalt-seawater interactions from 25 to 300°C and from 1 to 500 bar: an experimental study (abstract). EOS, Trans. Am. Geophys. Union 56, 1073. SEYFRIEDW. E. and BISCHOFFJ. L. (1977) Hydrothermal transport of heavy metals by seawater: the role of seawater/basalt ratio. Earth Planet. Sci. Left. 34, 71-77 SEYFRIED W. E. and Mom M. J. (1977) Origin of submarine metal-rich hydrothermal sotutions: experimental basalt-seawater interaction in a seawater-dominated system at 3OOC, 500 bar. Proc. of the Second Int. Symp. on Water-Rock Interaction, I.A.G.C., Strasbourg France, pp. IV 173-IV 180. Spoo~wl E. T. C. and FIFE W. S. (1973) Subsea floor metamorphism, heat and mass transfer. Contrib. Mineral. Petrol. 42, 287-304.

TAYLORH. P. JR (1974) The application of oxygen and hydrogen isotope studies to problems of hydrothermal alteration and ore deposition. Econ. Geol. 69, 843-883. THOMPXING. and MELSONW. G. (1972) The petrology of oceanic crust across fracture zones in the Atlantic Ocean: evidence of a new kind of sea floor spreading. J. Geol. 80, 526-538.

TOMA~~~NJ. and KRISTMANNSD~T~IR H. (1972) High temperature alteration minerals and thermal brines, Reykjanes, Iceland. Contrib. Mineral. Petrol. 36, 123134. WILLIAMSD. L., VON HERZEN R. P., SCLATERJ. G. and ANDERSONR. N. (1974) The Galapagos Spreading Center: lithospheric cooling and hydrothermal circulation. Geophys. J. Roy, Astron. Sot. 38, 587-608. WOLERYT. J. and SLEEPN. H. (1976) Hydtothermal circulation and geochemical Aux at mid-ocean ridges. J. Geol. 84, 249-275.