Alteration of the oceanic crust: Implications for geochemical cycles of lithium and boron

Gmhrmrca ef CosmcxhrmrcnActn Vol. 48. w. 557-569 Q Pcrgamon PressLtd. 1984. Pnnted in U.S.A.

GtR&-7037/84/$3.M)+ .I0

Alteration of the oceanic crust: Impli~tions for ge~hemical cycles of lithium and boron W. E, SEYFRIED, JR.‘, D. R. JANECKY’, and M. J. MOITL* ‘Department of Geology and Geophysics, University of Minnesota, Minneapolis, Minnesota 55455 2Department of Chemistry, Woods Hole Oceanographic Institution. Woods Hole, Massachusetts 02543 (Received July 19. 1983; accepted in revised form December 12, 1983) Abstract-Fresh tholeiitic basalt glass has been reacted with seawater at 15O”C, (water/rock mass ratio of IO), and fresh diabase has been reacted with a Na-K-Ca-Cl fluid at 375°C (water/rock mass ratios of I, 2, and 5) to understand better the role of temperature. basalt composition, and water/rock mass ratio on the direction and magnitude of B and Li exchange during basalt alteration. At 150°C slight but nevertheless significant amounts of B and Li were removed from seawater and inco~mt~ into a dominantIy smectite alteration phase. At 375”C, however, B and Li were leached from basalt. B behaved as a “soluble” element and attained concentrations in solution limited only by the B concentration in basalt and the water/rock mass ratio. Li, however, was less mobile. For example, at water/rock mass ratios of 1, 2, and 5, the percent of Li leached from basalt was 58, 70, and 92% respectively. This suggests some mineralogic control on Li mobility during hydrothe~~ alteration of basalt, especially at tow-water/rock mass ratios. In general, these results, as well as those for B, are consistent with the temperaturedependent chemistry of altered seafloor basalt and the chemistry of ridge crest hydrothermal fluids. Based on the distribution and chemistry of products of seafloor weathering, low (1;15OT) and hightemperature hydrothermal alteration of basalt, and the chemistry of ridge crest hydrothermal fluids, it was estimated that alteration of the oceanic crust is a Li source for seawater. This is not true for B, however, since the hot spring flux estimated for B is balanced by low-temperature basalt alteration. These data, coupled with B and Li flux estimates for other processes (e.g., continental weathering, clay mineral adsorption, authigenic silicate formation and formation of siliceous skeletal material) yield new insight into the B and Li geochemical cyctes. Calculations performed here indicate relatively good agreement between the magnitude of B and Li sources and sinks. The geochemical cycle of B, however, may be atiected by serpentinixation of mantle derived peridotite in oceanic fracture zones. Serpentinites are conspicuously enriched in B and if the B source for these rocks is seawater, then an additional B sink exists which must be integrated into the B geochemical cycle. However, until more data are available in terms of area1 extent of serpentinization, serpentite chemistry and isotopic composition, the importance of B in these rocks with respect to the 5 geochemical cycle remains speculative at best. INTRODUCTION

NUMEROUSINVESTtGATlONShave d~ument~ relatively large changes in the B and Li concentrations in weathered and hydrothermally altered basalt from the seafloor relative to the concentrations of these elements in unaltered equivalents of these rocks (MELSON et al., 1968; HART, 1970; THOMPSONand MELSON, 1970; THOMPSON,1973a; HUMPHRISand THOMPSON,1978; DONNELY et al., 1979; BERGERONet al., 1983). HART (1970), THOMPSON f 1973a), and DONNELLY el al. (1979), for example, have shown that B and Li are enriched in basalts altered by seawater at relatively low-temperatures. The opposite appears to be true, however, for hip-tem~rature, hydrothe~al alteration of basalt as evidenced by the chemistry of interior portions of seafloor greenstones (HUMPHRIS and THOMPSON,1978). Furthermore, mobility of Li during high-temperature alteration of the oceanic crust has been confirmed recently by the relatively Li-rich nature of ridge crest hydrothermal fluids emanating from vents on the Galapagos Spreading Center (EDMOND et al.. 1979a) and at 2 1“N, EPR (EDMONDet al., 1982). Analagous evidence for B, however, is less convincing (EDMOND et al., op. cit.), although a relatively small

positive B anomaly has been identified in hot-spring fluids at 21 “N, EPR (A. SPIVACK,personal commun.). Geochemical cycles of B and Li must be affected by processes related to ocean crust formation, aging, and reaction with seawater. Whether or not the net result of these processes contributes B and Li to seawater depends critically on the extent of high and lowtemperature alteration in the oceanic crust and the magnitude of B and Li exchange associated with these events. The purpose of this investigation is to document the direction and magnitude of B and Li exchange during relatively low and high-temperature hydrothermal alteration of basalt from experimental basaltseawater and ba~lt-~lution interaction at temperatures of 1SO’C and 375°C respectively. These temperatures were chosen to represent upper limits for low and high-temperature hydrothermal alteration of basalt at or near mid-ocean ridges. Low-temperature basalt-seawater interaction is typically characterized by a smectite dominated alteration assemblage, whereas high-temperature basalt-seawater (or seawater derived fluid) interaction appears to be responsible for the chemistry of ridge crest hydrothe~al kids (EDMOND et al., 1979, 1982) and the chemistry and mineralogy

557

558

W. E. Seyfiied, Jr.. D. R. Janecky and M. J. Mottl

of some greenstones and other metasomaticaliy altered rocks dredged from mid-ocean ridges (HUMPHRIS and THOMPSON, 1978; MELSON ef d., 1968). Characterization of the temperature dependent mechanism and magnitude of B and Li exchange during low and hightemperature hydrothermal alteration of basalt represents an essential ingredient to our knowledge of geochemical processes affecting the global balance of these two elements.

1.

Chemistry of starting experiments.

____-component

______ Na K

The 150”Cexperiment involved basalt glass-seawater interaction, whereas diabase and a Na-C&K-Cl fluid were used for the 37YCexperiments (Table 1). A Na-Ca-K-Cl fluid of the sort used here is consistent with our understanding of the compositional changes taking place in seawater chemistry as seawater reacts with increasingly greater amounts of basalt on descent to deeper and higher temperature portions of the submarine geothermal system (see discussion to follow and SEYFRIEDand Mo~L, 1982). Basalt glass and diabase were powdered to ~62 Ctm to enhance reaction rates. The 15O”Cexperiment was conducted at a water/rock mass ratio of 10 (500 bars) in a gold reaction cell, whereas the 375”Cexperiments were conducted at water/ rock mass ratios of I, 2, and 5 (400 bars), in a small volume titanium reaction cell. Experiments were performed using Dickson hydrothermal apparatus (DICKSON et al.. 1963; SEYFRIEDPI al., 1979). The advantage of using this apparatus is that it permits sampling of aqueous fluid at experimental conditions. Nine samples were taken from the lSO”C-experiment, and two samples were removed from each of the 375”C-experiments.

: SO4 CT* Ci

Basalt glass and diabase used for this study were identical to basalts used for experiments reported on previously (SEYFRIED and BISCHOFF,1979: Morr~, 1983a). These rocks were analyzed for B and Li by DC-Plasma Atomic Emission Spectroscopy. The method of analysis used for B combined sample digestion (0.5 g) by potassium carbonate fusion with separation of cations by cation exchange techniques (FLEEI. 1967). Exchange techniques were required to separate cations such as iron which cause spectral interference with B. Li concentration was determined on the dissolved salts from a low-temperature, multi-acid (HF-HCIO,) digestion using ag proximately I.0 g of rock. The Li and B concentrations obtained for the basalt glass and diabase are 7 C I ppm and 2.5 +- 0.5 ppm and 5.0 t 0.5 ppm and 5.7 + 0.5 ppm respectively (Table 1). These concentrations are typical of unaltered tholeiitic basalt from the seafloor. For example, HUMPHRISand THOMI?WN(1978 and references therein) reported Li concentrations of 2-R ppm for fresh oceanic tholeiitic basalts, whereas B concentrations were repofied as 3-8 ppm. A~JMENTO (1968) and AUMENTOand LONCAREW (I 969) reported B concentrations of 2-13 ppm for fresh basal& from the Mid-Atlantic Ridge at 45“N, while MELSONd al. (1968) reported B concentrations of
_--___‘GIESKES er al. (1982) analyzed the Li concentration In solution for the sample taken at 1968 hrs., and reported a value similar, but not identical to ours (Table 2). The difference between the two measurements, however. is within the error

.:r*ed",i

:.u~d

.._.__... __

;ppm

.._ 'L .,. j'

* (4 .'j,T

i

Basalt Class --2.5 2 0.5

Li

*Concentration

_~ __

10,320 383 393 1,240 2,620 2.37 18.510 '1.35 0.190

Rock Components E--

EL-perimental

fluids and rocks

--..- ._._--... -.. Na-K-C+*.‘ seawater (Ppm)

u LL

PROCEDURES

of our analysis.

Table

7.0 + 1.0

in mmoledkg

soluf:on.

Li and B concentrations in solution were also analyzed b? DC-Plasma Atomic Emission Spectroscopy. Standards were prepared using Copenhagen Seawater (chlorinity = 19.3675%) containing known additions of Li and B. Although precision was not formally estimated, duplicate analyses for B and Lr were usually within &2% at concentrations in solution near seawater values (4.35 ppm B; 0. I9 ppm Li). Duplicate analyses for Li at 0.05 ppm were typically reproducible to within 220% of the reported value. RESULTS At 375”C, diabase altered entirely to a mineral assemblage including: mixed layer smectite-chlorite, actinolite, and albite. and minor magnetite, pyrite, and chalcopyrite (SEYFRIEDand JANECKY,in prep). B entered solution in amounts limited only by the B concentration in the diabase and the water/ rock mass ratio (Table 2, Fig. 1). This was not true of Li. however, since the Li enrichment in solution for each experiment was less than that possible from fresh basalt (Table 2. Fig. I). This was especially true for experiments conducted at relatively low water/rock mass ratios and implies retention of Li by secondary alteration phases, The 150”C-experiment resulted in only pamal alteration of basalt glass to smectite and anhydrite (see SEYFRIEDand BISCHOFF.1979). and revealed an apparent “turnaround” in B and Li concentrations in solution-that is, after initially Increasing the concentrations of these species decreased (Table 2, Fig. 2). The Li decrease was gradual, and by the end of the experiment. Li concentration in solution decreased by approximately 2.5 times relative to the Li concentration in the starting seawater solution.’ The relative change in B COW centration in solution was less pronounced than for Ii. Nev.ertheless, B decreased from 4.35 1 0.10 ppm at the start to 3.8 +- 0. IO ppm at the end (Table 2). DISCUSSION

Nigh-lemperuture und Li exchange

basal&seawaler

interacrron- B

Li and especially B are leached from diabase and enter solution during high-temperature interaction with a Na-Ca-K-Cl fluid. In general, these results are consistent with those based on the chemistry of the interior portions of hydrothermally altered pillow basalts dredged from mid-ocean ridges (HUMPHRIS and THOMPSON, 1978; MELSON ef al., 1968), as well as with results of previously performed high-temperature. rock-water interaction experiments. l3.1 1s I i WO), ht

B and Li in the oceanic crust Table

2.

Concentrations of Li and B (ppm) in solution before and during interaction with basalt glass at 150%, water/rock ratio Of 10, and with diabase at 375% and water/rock ratio3 of I, 2, and 5.

15ow

IO

375OC

375%

375°C

SeaYateP

o/o* t/5 z/79 3/175 4/504 5/1344 6/1656 7/1968 B/2460 913564

4.35 4.35 4.40 4.43 4.51 4.50 4.35 3.80

0.19 0.19 0.19 0.28 0.31 0.16 0.11 0.07 0.07

5

w&ca-I(-c1

II/* 1,250 Z/1541

0.0 1.34 1.33

0.0 0.85 0.90

*

Na-Ca-K-Cl

010 IllOOO 2/2OCO

$206

0.0 1.68 -

O/O l/1040 z/2041

0.0 6.39 6.69

1

Na-Ca-K-Cl

-

0.0 2.75 2.01

(-) NM analyzed (*) Starting fluid composition

example, integrated results of numerous rock-water interaction experiments (ELLJS and MAHON, 1964, 1967; ELLIS, 1967, 1968, 1969) with chemical data from natural geothermal systems and concluded that B and Li are “soluble” elements during hydrothe~al alteration of igneous rocks and sediments. Soluble eiements are species present in hydrothermal fluids, which are controlled not by temperature and pressure dependent ~lution-minem1 equilibria, but primarily by the abundance of these components in the host rock and the effective water/rock mass ratio; that is, these elements are for the most part incompatible in alteration phases and partition effectively into solution with the dissolution of primary rock minerals. Thus, “soluble” elements can be usefully employed to identify the source of fluids in hydrothermal systems and the water/rock mass ratio effective during alteration. Clearly, from the results of our investi~tion, 3 can be defined as a “soluble” element, but Li cannot, since small but nevertheless significant amounts of Li are retained by hydrous alteration phases. At a water/rock mass ratio of 5, however, Li attains “soluble” behavior. Apparently, Li concentrations in solution are sufficiently low so as to preclude uptake by alteration mineral phases. K is even less “soluble” than Li at the conditions of the high-temperature experiments (Fig. I). Only about 22% of the K is leached from diabase at a water/ rock mass ratio of 1, while at water/rock mass ratios of 2 and 5,28 and 37% respectively of the K is removed to solution. There can be. little question that K concentration in solution is strongly affected by solutionmineral equilibria at the relatively high temperatures of the experiments (SEYFXJEDand JANECKY,in prep.). Application of experimental results to natural systems must be employed cautiously, since the experimental system is often different from the natural sys-

559

tern. This is clearly embodied in the concept of water/ rock mass ratio as applied to these two systems. As used here (experimental system). the water/rock mass ratio is simply the initial mass of fluid to rock in the reaction cell, whereas for the natural system, the water/ rock ratio is most often defined as the total mass of fluid which has interacted with the system, integrated through time, relative to the total mass of altered rock in the system (MOTTL, 1983b and references therein). The experimental system necessarily models a closed system, isothermal, batch process, while the natural system reflects reactions taking place in an open flowthrough system, between a fluid which is changing in temperature and composition and rock which is inhomogenously altered (MOTTL, 1983b). MOTTL (1983b) reviewed three principal mechanisms which can influence the chemical species in solution during basalt-solution interaction: (1) solutionmineral equilibria; (2) kinetics (reaction vs. flow rates); and, (3) depletion from the rock. Results of the present investigation imply that Li is influenced by solutionmineral equilibria. Thus, Li concentrations in solution, such as in ridge crest hydrothermal fluids, will reflect the P, T and com~sitional constraints of the system prior to ascent to the seafloor, and will not record previous alteration histories. B, however, being a “soluble” element will reflect the path the solution has followed through the system (Mont, 1983b). EDMOND et al. (1979a, 1979b) reported on the chemistry of fluids from submarine hot springs along the crest of the Galapagos Spreading Center. Although these fluids had entrained significant amounts of ambient seawater, it was still possible to resolve large compositional anomalies relative to seawater chemistry. For example, the Li concentrations in these hydrothermal fluids ranged from 0.194-0.4 1 ppm (ambient seawater contains approximately 0.17 ppm, CHOW and GOLDBERG, 1963). From this, and the hydrothermal fluid/ambient seawater mixing ratio, which could be assessed from Si/Temp. retations (EDMOND ez a/., 1979a), the Li content in the endmem~r fluid

FIG. 1. Ratio of B, Li and K concentrations in solution to that available in fresh diabase as a function of water/rock mass ratio. Experiments performed at 375°C 400 bars.

W. E. Seyfried. Jr., D. R. Janecky and M. J. Mottl

560

1.0

5.6 I

HOURS

32 I

HOURS

178 1000 5623 1 I I

1.0 I

5.6 !

32 1

178

1

1000

0.8

0.6

z? ‘(

0.4-

5623 1

-Limit

2

,

-I

of Li Solubillty

-_j

._I i

I

0.0

0.75

I

1

I

1.50

2.25

3.0

LOG

0.2-

1.50

TIME

2.25

3.0

3.75

(HOURS)

FIG. 2. Change in the concentrations of B and Li in solution during seawater-basalt glass mteracuorl JL 15O”C, 500 ba&, water/rock mass ratio of 10. (unmixed

hydrothermal fluid) was calculated as 7.9 to 4.75 ppm. These values suggest low water/rock mass

fluids issuing from the Galapagos vents (EDMOND cv al., 1979a). This is not altogether surprising considering

ratios for alteration at depth for the Galapagos geothermal system. This inference is consistent with the water/rock mass ratio computed by CRAIG ef al. ( 1980) from oxygen isotope systematics and measurements of total dissolved helium for Galapagos hot spring emanations. Furthermore, the Li concentration in hydrothermal fluid at 2 I “N, EPR is 5.65 ppm (EDMOND et al., 1982). This value agrees well with Galapagos data. Li concentrations in solution from our experiments are, in general, lower than those reported for endmember hydrothermal fluids at 2 I “N, EPR and Galapagos. This may reflect the existence of water/rock mass ratios for these hot spring systems lower than unity, or a basaltic substrate characterized by Li contents greater than that for the diabase used for our experiments. The former explanation is probably the more accurate of the two. For example. assuming a water/rock mass ratio of unity and considering the extent to which Li partitions between solution and altered rock at this water/rock ratio (Fig. I). then a Li concentration of 5.65 ppm in 2 1“N hot spring fluid requires a Li content of approximately 9.5 ppm for unaltered basaltic substrate, and this is considerably greater than the 2-8 ppm Li commonly reported for fresh tholeiitic basalt from the seafloor (HUMPHRIS and THOMPSON, 1978 and references therein). Thus, a water/rock mass ratio less than unity may be necessary to account for the relatively high Li concentration in ridge crest hydrothermal fluids. This is especially true for the Galapagos hydrothermal system. A B anomaly was not observed in hydrothermal

the magnitude to which the endmember hydrothermal fluid mixed with ambient seawater which contains a relatively high B concentration (4.6 ppm). However. hydrothermal fluids at 2 1“N, EPR do reveal a slight positive B anomaly (A. SPIVACK,personal commun. ), The magnitude of this anomaly for the endmember hydrothermal fluid is 1.05 ppm. Taken at fact value. and considering the “soluble” nature of B, this Implies either the existence of B-poor source rocks at 2I”N. or a process more complex than simple leaching of B from basalt during high-temperature interaction with seawater. Another possibility, of course, would be aiteration at a high water/rock mass ratio. but. as noted previously this is inconsistent with a host of other geochemical data for the 2 1ON hydrothermal system SPOONERand FYFE (1973), LISTER ( 1973, I9743. and WOLERY (1978) suggested that mid-ocean ridge hydrothermal systems are recharged by geographically dispersed limbs of cold seawater and discharged by relatively rapid ascent of hot fluid through localized fracture systems. Detailed studies of heat 80% at the Galapagos Spreading Center (GREEN et ~11..1% 1) and at the Juan de Fuca Ridge (DAVIS er a/., 1980) demonstrate the existence of high and low heat flow. These heat flow patterns are best explained by the presence of upwelling and downwelling limbs of hydrothermal circulation cells. As noted previously. however, it is probably inaccurate to assume that the recharge fluid of the downwelling limb is normal seawater. especially those tluids attaining sufficiently great depths and cotrespondingly high enough temperatures to generate EPR. 2 1“N-type hydrothermal exhalations. C”hcmica!

B and Li in the oceanic crust reaction with basalt and possibly sediment at reIativeIy low-temperatures must occur and to some degree, modify the composition of this fluid rehttive to that of normal seawater. For example, SO; and Mg” are probably lost and Ca++ gained by these fluids owing to a combination of alteration processes in the upper regions of the submarine geothermal system. These reactions, especially that ofS0~ removal from solution, are consistent with changes in seawater chemistry observed in basalt-seawuter experiments at temperatures as low as 150°C (SEYFRIEDand BISCHOFF,1979), and with the chemistry and mineralogy of met&a&s from mid-ocean ridges. The experiments show that SO; and Mg++ are rapidly and entirely removed from solution owing to anhydrite and smectite formation respectively. To remove SO: qua~ti~tively from solution, however, requires addition of Ca”, and this is a~ornp~~ed by exchange of basalt-Ca for seawater Mg+‘. If SO: were available in solution or as anhydrite during greenschist facies alteration of basalt leading to for* mation of seafloor metabasahs, then these rocks would be appreciably more oxidized than they actually are. For example, basalt-seawater interaction experiments at temperatures as low as 200°C (MOTTL et al., 1979) provide unequivocal evidence for SO; reduction. Thus, the general lack of oxidation (HIJMPHRISand THOMPSON, 1978) and scarcity of hematite in seafloor metabasalts indicate that little seafloor SO; was availabbfeduring hydrothe~~ aIteration. Another modification of seawater chemistry on descent into the submarine geothermal system may involve B removal. That this couM occur is consistent with the time dependent change in 3 concentration in solution observed during our 150°~-~x~~rnent (Fig. i)Yas well as with results of chemical studies of alteration products of low-temperature, basalt-seawater interaction (f)ONNELLYet al., 1979) and the conspicuously high B conc~ntmtio~s in the exterior portions of greenstones (HUMPHRIS and THOMPSON, 1978; MELSON ef al., 1968; VALLANCE, 1974). Thus, it is possible that downwelling, ridge crest hydrothe~a~ fluids are characterized by B c~n~ntmtions lower than that of normal seawater. Subsequent reaction with basalt at high-temperatures, however, would result in B enactment in solution in amounts regulated by the effective water/rock mass ratio during alteration and the B concentration in basalt. A mechanism such as this provides an explanation for the relatively low ap parent anomaly for B in 21 “N, EPR hydro~e~al fluids. The actual B anomaly may be as great as the total B concentration in these fluids. Only by careful study of the B isotopic systematics of ridge crest hydrothermal fluids can the two B sources be adequateiy resolved. Whether or not afteration history affects the magnitude of an anomaly for a particular component in ridge crest hydr~the~al fluid refative to seawater chemistry, depends on: f 1) the tendency of the species in question to participate in relatively low-temperature alteration processes; (2) element mobility during high-

561

temperature bait-~iution jntemction; and, f 3) the ratio of eXement concentration in basalt to that in seawater. Since B is effectively extracted from solution at relatively low-temperatures, extracted from basalt at hip-tem~mtu~ and occurs in seawater and basalt in nearly equal amounts, its concentration in ridge crest hydrothermal fluid cannot be used to assess unamb~guousiy deep-seated, h~~-t~rn~~ture alteration processes without carefui consid~mtion of alteration history. Li, like B presumably participates in secondary mineral formation during low-temperature basah-seawater interaction (DONNELLYet al., 1979). Hence the Li conc~ntra~on in solution may actually be less than in seawater prior to deepseated, high-temperature hydrothermal alteration of basah/diabase. Unlike B, however, the Li conc~ntmtion in seawater is significandy less than in b~tjdia~, and since Li is relatively, although not entirely, mobile during hightemperature basalt-solution interaction, ridge crest hydrovers fluid will be characterized by a large positive anomaly. The magnitude of the Li anomaly in endmember hydrothermal fluid can be used to compute, for example, water/rock ratio during hydrothermal alteration at the last point of equilibration between solution and rock provided experimentahy determined partition coefficients (see Fig. 1) are available for appropriate conditions. B and ti$ux us~~~ate~ with Hugh-temperature f>lSO”C} alteration ofthe ocearric crust-hot spring jlux &NKwS et a!. (19’78) calculated the convective hydrothe~al heat flux to be 4.9 I: 1.2 X IO’” Cal/year. This estimate was based on the 3He-temperature anomalies for fluids emanating from Galapagos hot springs and the rate of jHe loss from the atmosphere. Int~stin~y, a similar estimate ( - 5.4 X 101’ caI/yearf was arrived at based on the discrepancy between the theoretical heat production associated with seafloor spreading and heat flow observed at mid-ocean ridges (SLEEPand WOLERY, 1978). Although similar in magnitude, these estimates yield very different values for the flux of seawater through mid-ocean ridges because the heat flux caiculated by SLEEPand WOLERY( 1978) includes that portion of the oceanic crust from the ridge crest to the palm where the observed and theoretically calculated heat flow merge. Theoretical and observed heat tlow curves for mid-ocean ridges merge at 15 to 70 m.y. (AN~E~~N et al., 1977). However, the flux estimated by JENKINSet ai. ( 1978) is applicable only to the axial zone of mid-ocean ridges and necessarily neglects off-ridge hydro~be~~ activity. The conspicuous ‘He anomaly at a water depth of 2000 m in the South Pacific (EDMONDet al., 1982: CRAIG et al., 1975) is difficult to explain without appealing to a ‘He source near the crest of the East Pacific Rise. Thus, the agreement in the magnitude of the heat now anomaly calculated by these two independent methods is puzzling and possibly reflects non-steady state pro-

W. E. Seyfiied, Jr.. D. R. Janecky and M. J. Mottl

562

cesses associated with 3He-temperature systematics of ridge crest hydrothermal activity, or perhaps the existence of high 3He, low-temperature fluid emanations very near ridge axes (S. R. HART. pers. commun.). To date emanations of this sort have not been recognized. The hydrothermal heat flux at the crest of midocean ridges proposed by SLEEP et al ( 1980) is 3 X IO* Cal/cm’ (seafloor). Since about 3 km’ of new crust is created on all plates each year by seafloor spreading (CHASE, 1972), the ridge crest heat flow anomaly is -9 X 10” Cal/year, and the mass of water (Q) circulating through the crust is approximately:

Q=

9 X 10’” ca]/yr Cw(T, - T,)

= 2.35 f IO’” g/yr

where Cw is the average heat capacity of water (- 1.I cal/g”C. from HELGESON and KIRKHAM. 1974). T, = 2”C, and Tz = 350°C. integrating this flow rate with the average Li anomaly (6.8 ppm) measured for Galapagos and 2 1“N, EPR hydrothermal fluids (EDMOND d al., 1979a; 1982) yields a hot spring tlux for Li of 16 X 10”’ g/yr. The B hot spring flux is more difficult to estimate because of our uncertainty of the relative importance of seawater-B and basalt-B in ridge crest hydrothermal fluids. To obtain a maximum estimate. however, we will assume that B dissolved (5.5 ppm B) in 21 “N, EPR hot spring fluid is entirely basalt derived. This yields a hotspring flux for B of 13 X IO’” g/yr. Flux estimates for B and Li computed here are significant. The Li hot spring flux is approximately I .7 times the Li river flux, whereas for B, the hot spring flux is 0.27 times the B river flux (Table 3). Implicit in these flux calculations is the assumption that the temperature of end-member hydrothermal fluid and concentrations of components in this fluid represent steady state values. This assumption, however, is tenuous in light of our limited data base, and the now established episodic nature of ridge crest hydrothermal activity (MACDONALDet al.. 1980; MOTTL, 1983a). Low-temperature basalt-seawater mtc~ruction-B and Li exchange THOMPSON and MELSON (1970) documented enrichment of B and Li in basalt weathered at seafloor

conditions. These data are consistent \nth B and Lr abundances in the most highly altered basalts from DSDP Leg 51 (DONNELLYet ul.. 1979). However, in sharp contrast with basalt weathering on the seafloor. which appears to become increasingly altered up to 57 m.y. (THOMPSON, 1983: LAWRENCEand GIESKES. I98 1) basalt alteration recovered during ocean crusta] drilling, such as that from DSDP Sites 332B. 417A and 4 l8A. has formed very early in the history of the ocean crust. For example, HART and STAIJDIGEL (1978) and RICHARDSONer al. (1980). using Rb/Sr isochron ages of vein smectite-palagonite and celadonite and *‘Sr/*%r systematics of vein calcites, have convincingly demonstrated that mineralization occurs within 5-10 m.y. of the age of formation of the crust. Mineralization may be extended back to approximately 15 m.y. after crust formation as suggested by STAl:DIGEL (pers.commun.), based on recent revisions o! the Cretaceous time scale, thereby more correctly establishing the age of the MO magnetic anomaly and age of ocean crust at 417A. Nevertheless, 81 is clear that the duration of vein mineralization in upper sections of oceanic crust is relatively short lived and mosl probably associated with seawater circulation and recharge of deep-seated hydrothermal systems near actively spreading mid-ocean ridges. This is consistent with heat flow circulation models of SLEEPand War ERY (1978) and ANDERSONet al ( 1977). Thus, temperatures of vein formation and mineralization for DSDP alteration may have exceeded that of ambient seawater, and alteration temperatures a.~high d.s I JO170°C may have been attained by some hydrothermal fluids in deeper portions of layer II of the oceanic crust (STAKES and O’NEIL, 1982). Even though these alteration temperatures are relatively high (zeolite facies). they are not high enough to cause elimination of smectite, especially saponite (STAKES and O’NEII . 1982) as a dominant alteration phase nor to preclude B and Li removal from seawater as evidenced hv results o] the present investigation. In addition to temperature. however. alteration phase chemistry probably plays a critical role in determining whether or not B and Li are removed from solution. For example, since Mg and Li have similar ionic radii (-0.68 A), it may be necessary thal some Mg be contained in an alteration phase befort- Li re’ moval from solution can occur. This interpretation 1s consistent with the time dependent changes in Mg and Li concentrations in solution during the 150°C experiment (Fig. 3), which showed that Mg removal from solution preceded that of Li. Mg removal was associated with formation of Mg(OH), and/or Mg$i,O,,,(OH), layers in a smectite alteration phase Furthermore, the decrease in B concentration in solution may also be related to the Mg-rich chemiste ofthe smectite alteration phase. RH~ADESPI ~ri i 19701 reported that significant amounts of B can be adsorbed onto Mg(OH)2 surfaces. Alternatively, B removal from solution may be accounted for by isomorphic, rcplaccment of tetrahedrally coordinated Si and A] 11-1tl-rc

B and Li in the oceanic crust

563

“normal” oceanic crust, and integrated with appropriate B and Li enrichment factors for these phases, 40 7 then a first approximation can be made to assess the effects of low-temperature alteration on B and Li geochemical cycles. Basalt and basalt alteration recovered 20 6 during DSDP Leg 52 and 53, Site 4 18A appears to be 2 ideally suited for this purpose because the forma1 age E-z 0 5 PH ai of crust formation at this locality is approximately I IO m.y. B.P. (STAUDIGELef al., 1979) and thus these 5 -20 i4 rocks contain the combined effects of seafloor weathering and low-temperature hydrothermal alteration. -40 3 JOHNSON (1979) described in detail the dist~bution of various alteration phases in basalt recovered from Site 4 18A. Drilling at this site penetrated approximately 0 6 12 18 24 30 36 550 m of basalt and alteration is characterized by an TIME IN HOURS X 100 assemblage analogous to that from other DSDP sites (e.g.,332B, MUEHLENBACHS,1977:3968.M~~~~~~FIG. 3. Change in pii and concentration of Li and Mg in solution during seawater reaction with basalt glass at 15O’C. BACHSand HODGES, 1978). Considerably less alteration Mg and pH data are from SEYFRIED and BISCHOFF (1979). exists at this locality than at nearby Site 4 I7A. SmectiteWed veins account for about 7. I % of the core, while carbonate veins account for 5%. These abundance essmectite. Numerous investigators studying adsorption timates, however, were based on recovered core only processes and diagenetic modifications involving clay (72% ofthe drillhole). HART and STAUDIGEL(1982) minerals have documented just this sort of reaction and STAUDIGELand HART (1983) determined indi(FLEET,1965; HARDER,1970). rectly the portion of secondary phases in the remaining 28% of the core and added these to the distribution given by JOHNSON (1979). These results indicate that 3 and Li jlux associated with tow-temperature secondary alteration at DSDP Site 418A consists of: (< f.W’C) alterati~tt r$the oceanic crust 7% carbonates; 10% smectites; and 16% palagonite. The importance of relatively low-temperature, baThe remaining 67% of the core consists of unaltered salt-seawater interaction to the geochemical cycles of glass and variably altered breccia material, pillows and B and Li is dependent on the amount of this type of massive flows. Using these data and the scheme of alteration in the oceanic crust. Results from DSDP HART and STAVDIGEL (1982)and STAUDKXL and reveal, however, that alteration is heterogeneously HART (1983), we can sub-divide bulk alteration at Site distributed in oceanic rocks because of inherent vari4 18A into smectite-palagonite (26%) and whole-rock ations in permeability, porosity, and seafloor topog(74%) alteration. raphy. These features influence the extent ofseawaterB and Li concentrations for smectite-palagonite albasalt interaction (MUEHLENBACHS, 1979 and referteration can be estimated from data from Site 41’7A ences therein). For example, basalt from Leg 5 I, hole (DONNELLY erai.. 1979). As noted previously. rocks 4 I 7A is very nearly, if not entirely. altered and replaced recovered from this site are characterized by an abunby a suite of secondary phases including protoceladance of highly altered material dominated by various donite. smectite (montmorillonite and saponitef. calsmectite-rich phases. B and Li concentrations for this cite, K-feldspar, various zeolites (SCHEIDEGGERand material aregivenby DONNELLY etal. (1979)as 110 STAKES,l979), and hematite (HUMPHRISC~al.. 1979). and 55 ppm respectively. Assuming these concentra~o~espondingly, these rocks contain strikingly high tions apply as well to smectite-palagonite alteration B. Li, KZO, and H20, low CaO, and as the presence from Site 418A, and integrating this with B and Li of hematite indicates, are extremely oxidized (DON- concentrations for whole rock alteration from 4 I8A NELLY c'f al., 1979). However, that site 417A is char(STAUDIGELet al., 1979) yields B and Li fluxes for acterized by a relatively thin sedimentary sequence low-tem~rature hydrothe~al alteration and weathsuggests that the underlying basalt may have been a ering of the oceanic crust of 16-19 X 10” g/yr and positive topographic feature on the seafloor shortly 4.4-7.4 X IO” g/yr respectively (Table 4). The range after crust formation (DONNELLY L'I al., 1979). This for these flux estimates chiefly reflects our uncertainty would facilitate seawater circulation and account for with regard to Band Li concentrations in fresh oceanic the extremely altered nature of these rocks. It is for thoteiitic basalt. this reason, that the magnitude of chemical exchange Flux estimates for B and Li computed here are in between seawater and basalt observed here cannot he reasonable agreement with those proposed by THOMPextrapolated, a priori, on a global scale, since the SON (I 983) considering the great number of unceroceanic crust as a whole probably does not experience tainties and assumptions involved in calculations of this sort. For example, THOMPSON (ibid) estimated B such pervasive alteration. If. however, the distribution of alteration phases can be estimated for a section of and Li fluxes for low-temperature basalt-seawater re60

8

GAIN

I

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U t. Se~tned. Jr.. I). K. Janecky and M. J. Mottl

actions of 8.3 X 1O’Og/yr and 6.6 i: IO”’ g/yr respectively. These flux estimates, as well as others assessed by THOMPSON (I 983). represent the combined effects of seafloor weathering, low-temperature alteration in deep portions of oceanic basement and off-ridge. lowtemperature hydrothermal activity. Results from DSDP Site 418A were used to assess deep basement alteration, the chemistry of dredged basalt used to compute the effects of seafloor weathering. while the extremely altered rocks characterizing Site 4 17A were evoked to assess off-ridge, low-temperature hydrothermal activity. However. in light of the similar temperatures of alteration at Sites 4 I 7A. 4 I7D and 4 1XA (LAWRENCE, 1979: MUEHLENBACHS, 1979). and that 4 I7A represents a topographic basement high, it may not be appropriate to define uniquely chemical fluxes for an off-ridge hydrothermal component from Site 4 17A alteration. Comparison of fluxes associated with ridge crest hydrothermal activity (hot spring tlux), and iow-temperature basalt alteration (Tables 3 and 4) indicates that alteration of the oceanic crust is a source of LI for seawater. This is not true of B. however. since the B hot spring flux is more than compensated for by low-temperature basalt alteration. Nevertheless. additional processes must be evoked to offset B and Li contributions to seawater from continental weathering and the Li hot spring flux, if the concentrations of these species in seawater are to be maintained at steady state levels. Geochemical cycle of boron The dominant mechanism by which B is removed from seawater appears to be adsorption by continentally derived clay minerals. That clay minerals. especially iliite and smectite, adsorb B from seawater has been documented by numerous investigators (GOLD, SCHMIDTand PETERS, 1932; KEITH and DEGENS, 1959; SHAW and BUGRY, 1966; REYNOLDS, 1965; OHRDORF, 1968). For B, HARRISS (1969) estimated that clay mineral adsorption reactions remove approximately 33

i IO”’ g/yr B from seawater, which IS nearit 83% o! the river input of this species (Table 5). Other removal processes noted by HARRIS (of’. N.) include authigemc silicate formation (6 X IO”’ g/yr or 15% of the rive; flux), and formation of siliceous skeletal material (! * IO”’ g/yr or 7.5% of the river flux). The sum of the B removed from seawater by these processes IS. tberefore, 42 X IO”’ g/yr, and when added to the B flux estimated for low-temperature alteration of the oceanic crust ( 16- 19 X IO”’ g/yr). a value of 5X-h f IO”’ gi yr can be computed for the total amount of B removed from seawater annually (Table 5). This value IS IIS reasonable agreement with the input flux or B from rivers and ridge crest hydrothermal activity fli‘abie i j However, B may be involved in an additrctnal geochemical process. THOMPSON and MELSON (I 970) have observed that oceanic serpentinites sampled in dredge hauls are conspicuously enriched in B, containing approximately 70-100 ppm B. whereas the unaltered equivalents ol these rocks contain 110 ppm B. ‘The source :rf B fog serpentinites is most likely seawater (WENNER and TAYLOR, 1973; SAKAI and TSU-IWMI. 1978 1. BoNA?‘~ I t? al. ( 1980) has proposed seawater-peridotne Inter.. action to account for carbon, oxygen, and strontium isotope ratios of aragonite veins m oceanic serpentm ites. SEYFRIED and DIBBLE (1980). JANECKI and &IFRIED ( 1980), and JANECKY ( 1982) have documented B removal from solution during seawater-peridotite experiments at 200-3OO”C, 500 bars. Thesr studier, showed that the direction and magnitude of B exchange is extremely sensitive to temperature and soiutron pti. For example, neutral to moderately acid solutions of' seawater origin did not lose B during reactron with peridotite and serpentinization at a temperature ot 300°C (SEYFRIED and DIBBLE, 1980: JANWK~ . 1982). However for a neutral to alkaline solution at 200°C.

B and Li in the oceanic crust

565

and chemistry and/or a different chemistry for “pore” or for distinctly alkaline solutions at 300°C B decreased in solution during serpentinization (JANECKY fluids which infiltrate these rocks. Regardless of the exact mechanism by which B is removed from seawater and SEYFRIED, 1980; JANECKY, 1982). This is especially obvious during quench because solution pH rises and enriched in serpentinized peridotite (i.e., seafloor weathering or hydrothermal processes), the end result (due to retrograde reaction) as temperature decreases. is a sink for B of potential importance to the B geoThus B removal from seawater is possible during serchemical cycle. However, until more data are available pentinization provided an appropriate physiochemical on the distribution, chemistry and possibly B isotopic environment is maintained. composition of serpentinites in the oceanic crust, esOceanic serpentinites have been recovered fretimates of the relative importance of formation of these quently from dredge hauls of the walls of large fracture rocks to the geochemical cycle of B will be speculative. zones over wide areas of the north and equatorial MidAtlantic Ridge, and the Mid-Indian Ridge, and are often associated with basalt, gabbro and the metaGeochemical cycle of lithium morphic equivalents of these rocks (QUON and EHLERS, 1963; NICHOLU et al., 1964; LE PICHON, 1969;MELIdentifiable sources of Li for seawater include that SON et al., 1968; THOMPSONand MELSON, 1972; BONprovided by the dissolved load of rivers (9.4 X 10” A-II, 1976; BONA-I-~Iand HONNOREZ, 1976; Fox et g/yr) and the hot spring flux (- 16 X 10” g/yr) at al., 1973; BONATTIand HAMLYN,1978). These fracture mid-ocean ridges (Table 3). Assuming steady state, the zones are characterized by transverse valleys and ridges, Li influx must be balanced by appropriate removal and expose enormous sections of oceanic crust and mechanisms. Certainly, low-temperature alteration of upper mantle rocks. For example, the south wall of basalt comes into play here. As previously estimated, the Vema Fracture Zone reveals a section of over 4 this process removes only 4.4-7.4 X 10” g,/yr Li (Table km composed entirely of serpentinized peridotite 4). Thus, like B, an additional mechanism must be (B~NATTI and HONNOREZ, 1976). Interestingly, fraccalled upon to account for the apparent Li imbalance. ture zones offsetting the East Pacific Rise are not charLi enrichment is associated with pelagic marine clays acterized by transverse ridges nor outcrops of ultrarelative to Li contents of their non-marine countermafic rocks. Apparently, rate of seafloor spreading inparts. For example, KEITH and DEGENS ( 1959) refluences significantly tectonic forces responsible for ported the Li content of Pennsylvanian-age marine vertical protrusions of mantle derived igneous rocks and non-marine shales to be 159 ppm and 92 ppm (DICK et al., 198 1). Consistent with this interpretation respectively. These investigations reported the same is the relative abundance of serpentinized ultramafic relative change in Li contents for recent marine and rocks in fracture zones along the very slow spreading non-marine muds from Hawaii. OHRDORF ( 1968) recircum-Antarctic plate boundary. ported a change of only -45 ppm between the Li It appears from these investigations that a significant contents of marine and non-marine carboniferous portion of the oceanic crust may be serpentinized peshales from Germany, whereas HOLLANDet al. ( 1976) ridotite. Integration of this observation with the fact reported a difference of approximately 20 ppm in the that oceanic serpentine phases contain appreciable B, Li contents of marine pelagic clay (50 f 5 ppm) and which can be supplied by seawater as demonstrated the average igneous rock (30 + 5 ppm). Thus ample by experiment, leads one to conclude that the process evidence exists to support the thesis that Li is extracted of serpentinization may be an important mechanism from seawater by continentally derived pelagic clays for B removal from seawater. For example, if serpenand this presumably represents the additional Li sink. tinized peridotite constitutes just 1% of newly formed The accuracy of the relative difference in Li contents oceanic crust (12.3 km3/yr; FYFE and LONSDALE,198 1) between what are inferred to be “marine” and “nonand is characterized by a B enrichment of 70 ppm, marine” shales is predicated on the validity of a host then this would account for the B imbalance noted in of paleoenvironmental factors. Furthermore, this proTable 5. cedure implicitly assumes that shale chemistry accuIt is necessary to point out, however, that the ap rately records the chemistry of clayey sediments at the time of sedimentation. However, this is inconsistent parent B enrichment in serpentinized peridotite may with results from Gulf Coast drill cores which show not reflect the hydrothermal process resulting in the formation ofthese rock types, but rather it may indicate that post-depositional effects of temperature and pressure induce large compositional and mineralogic a relatively late-stage chemical modification to a previously serpentinized rock owing to reaction with sea- changes in sediments (HOWER et al., 1976). For these reasons, the technique employed by HOLLANDef at. water at low-temperatures, that is seafloor weathering. (1976) to document the magnitude of Li adsorption A modification of this sort to serpentinite chemistry would be similar to that observed for the margins of by marine pelagic clay is most useful provided the chemical data used to construct the average analyses greenstones dredged from relatively young ocean crust (HUMPHRIS and THOMPSON, 1978). The greater B of marine clays and igneous rocks are truly representative. concentration of oceanic serpentinites relative to the If we assume that continentally derived clays extract exterior of greenstones may reflect different times of approximately 20 ppm Li from seawater (HOLLAND exposure to seawater, different whole-rock mineralogy

W. E. Seyfried, Jr.. D. K. Janecky and M. J. Mottl

566

et al.. 1976), then by knowing the rate at which rivers discharge these phases to the oceans, we can estimate the amount of Li removal from seawater by adsorption processes. The total amount of suspended sediment delivered to the oceans annually by rivers is 183 X 10“’ g (HOLEMAN, 1968; GARRELS and MACKENZIE, 197 1). As-

suming the suspended sediment load of rivers is equal to the average shale plus sandstone, in a ratio of 87: I3 (GARRELS and MACKENZIE 197 1). and that the clay mineral content of the average shale is 61% (SHAW and WEAVER, 1965), the clay mineral content of the suspended load of rivers can be estimated to be 53%; that is, 9.7 I X lOI g clay are contributed to the ocean annually by rivers. Considering the 20 ppm enrichment factor with the amount of clay discharged to the oceans from rivers (9.71 X lOI g/yr), yields a value of 19.4 X 10” g/yr for the amount of Li extracted from seawater by clay mineral adsorption. This value provides an overestimate of Li adsorption by marine clays since all continentally derived clays do not have the opportunity to react with seawater. Nevertheless, integrating this with other Li sources and sinks results in a good overall balance for processes considered here as constituting the Li geochemical cycle (Table 5). This may be fortuitous, however, considering that we did not assess the role of marine organisms or authigenic silicate formation in marine sediments on Li uptake from seawater. Clearly, these processes must influence to some degree the Li geochemical

cycle, as they do that

of B, but insufficient data exists to evaluate them unambiguously at this time. It is our opinion, however, that if these processes are found to be significant. then they must be included in the Li geochemical cycle at the expense of our estimated flux for Li adsorption by continentally derived clay minerals. The Li river flux, hot spring flux and flux associated with low-tem-

perature alteration of the oceanic crust are reasonably well constrained and most probably will not change greatly with further examination. SUMMARY Experimental results have shown B and Li to be sensitive indicators of temperature during interaction of basalt glass and diabase with aqueous solutions. B is quantitatively leached from diabase at 375°C and attains concentrations in solution limited only by the B concentration in diabase and the water/rock mass ratio emve during alteration. In contrast, only about 60% of basalt-derived Li enters solution during hightemDerature alteration at a water/rock mass ratio of 1. At water/rock mass ratios of 2 and 5, greater amounts of Li are leached. At a water/rock mass ratio of 5 and 375’C, for example, approximately 92% of the Li available in basalt enters solution. At lSO”C, however. B and Li are removed from solution during basalt

alteration. In general, experimental data are in good agreement with the chemistry of seafloor weathered and high-temperature, hydrothermally altered basalt.

which typically reveal B and Li enrichment and de,, pletion respectively, relative to the concentrations of these species in inferred unaltered precursors. The chemistry of ridge crest hydrothermal fluid 13 influenced both by solution-mineral equilibria and bq total transfer of basalt components to solution. This

latter aspect effectively defines “soluble” elements which can be usefully employed to estimate the water/ rock mass ratio of deep-seated alteration processes. However, if the concentration in solution of a particular component is to be used for this purpose, then thr chemical nature and evolution of the solution involved in the hydrothermal episode must be known Since seawater undoubtedly represents the “starting” composition for solution descending into the submarine geothermal system, modification of seawater chemistr) owing to formation of temperature-dependent alteration phases may influence the magnitude of chemical exchange between basalt and solution as inferred from the chemistry of ridge crest hydrothermal fluids. Commonly it is assumed that a fluid of seawatm chemistry is implicitly involved in the event resulting in formation of end-member hydrothermal fluids. This may not be the case, however, especially for seawater components exhibiting a tendency to partition into low-temperature alteration phases. This may be an important process affecting the apparent magnitude of the observed B anomaly in ridge crest hydrothermal fluids. This is of lesser importance to the Li concentration in hot spring fluids because of the relatively small Li concentration in seawater and the tendency of Li to be controlled by solution-mineral equilibria at high temperatures. Comparison of fluxes associated with ridge crest hydrothermal activity (hot spring flux) and low-temperature basalt alteration indicates that alteration of the oceanic crust is a source of Li for seawater. This IS not true of B, however,

since the B hot spring flux 1s

effectively offset by low-temperature basalt alteration. Integrating B and Li fluxes associated with low and high-temperature basalt interaction with the river fluxes of these species indicates that other geochemical process must play a role in maintaining B and Li at steady state concentrations in seawater. Adsorption onto continentally derived clay minerals is an important sink for both B and Li. Additional sinks for H include: siliceous test formation by marine organisms; authigenic silicate formation; and, possibly serpentinization of mantle derived peridotite.

Acknowledgments-We wish to thank Drs. K. Muehlenbachs. H. Staudigel and S. R. Hart for their thoughtful critiques which greatly improved this paper. We would also like to thank I%. Ci. Thompson for providing us with a preprint of his recent paper on basalt-seawater interactions. These experiments were performed in the hydrothermal and analytical laboratories at the University of Minnesota with funding from NSF Grant OCE-8018644. Funding for M. J. Mottl was orevided by NSF Grant OCE-8110913. This is contribution number 1056 of the School of Earth Sciences, Departmem of Geology and Geophysics, Universily of Minnesota, and Woods Hole Oceanographic Institution contribution numb

5522.

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