Gmchrmm cr Cusmochwnica Acla Vol. 47. pp. 1613-1624 0 Pergamon Press Ltd. 1983. Printed in U.S.A

0016-7037/83/091613-12$03.00/0

Alteration of oceanic crust and geologic cycling of chlorine and water

EMI ITO,* DAVID M. HARRIS** and ALFRED T. ANDERSON, JR.~

*Department of Geology and Geophysics University of Minnesota 3 10 Pillsbury Dr., SE. Minneapolis, MN 55455 **Department of Geology University of Alberta Edmonton, Alberta Canada, T6G 2E3

TDepartment of Geophysical Sciences The University of Chicago 5734 S. Ellis Ave. Chicago, IL 60637

(Received July 20, 198 1; accepted in revised f&-m June 14, 1983)

Abstract-We report new estimates of transport rates for Hz0 and Cl between the mantle and surface reservoirs. Our estimates take into consideration alteration of oceanic crust. especially that of plutonic rocks, and possible subduction of sediments. The effect of (hydrothermal) alteration on the Cl budget seems to be negligible, but the effect on the Hz0 budget is significant. Altered oceanic crust (excluding sediments) contains about IO times as much Hz0 as the unaltered crust, and its subduction may result in a net transport of H20 to the upper mantle in subduction zones. However, the rate of expulsion of Hz0 from the mantle by subduction-zone magmatism is comparable to the amount released by ridge magmatism, and is only about 10% of the amount subducted. Therefore, about 90% of the subducted Hz0 must be returned to the mantle or returned to the crust by other processes. In addition, subduction of oceanic sediments to mantle depths will result in (1) a further increase in the return rate of H*O to the mantle reservoir, and (2) possible net transfer of Cl to the mantle, depending on the rate of pore water expulsion.

INTRODUaION

THE RATES at which the substances such as H20, COZ, Cl, and N2 were accumulated in the surface

reservoirs (e.g., air, oceans, and sedimentary rocks)

provide important clues to the way the Earth formed and how it works (GOLDSCHMIDT, 1933, 1954;

BROWN, 1949; RUBEY, 195 1; ANDERSON, 1974. 1975; SCHILLING et al., 1978). RUBEY (1951) dem- onstrated that seawater and sedimentary rock contain more Hz0 and Cl than can be derived from weath- ering of igneous rocks. He proposed that outgassing

of magmas by volcanism and crystallization gradu- ally supplied the surficial Hz0 and Cl during the his-

tory of the Earth. Modern plate tectonic theory suggests that present-

day subduction of altered oceanic crust results in a recycling or return of Hz0 and Cl from the surface reservoir (atmosphere, hydrosphere, and crust) to the mantle and back again. Consequently, subduction can change the net rates of outgassing. Our aim is to

review the effect of plate tectonic processes on the movement of HZ0 and Cl between the mantle and surface reservoirs incorporating results of a study on the alteration of oceanic crustal layer 3 (ITO and AN- DERSON, 1983) and direct measurements of Hz0 and Cl in the quenched magmatic liquids from Hawaii, mid-ocean ridges and subduction zones (HARRIS, 1981; HARRIS and ANDERSON, 1983a,b; HARRIS, unpub. data).

ANDERSON (1974, 1975) estimated that arc mag- matism alone may have released enough Hz0 and Cl in 1 to 10 Ga to account for the present quantities in the surface reservoir. Anderson further estimates that arc volcanism releases 10 times more HZ0 than, and about the same amount of Cl as, the oceanic ridge magmatism. He suggested that Hz0 and Cl in igneous rocks produced at the ocean ridges may be

refluxed during subduction-zone magmatism, thereby transporting HZ0 and Cl to the surface reservoir in a two-stage process without need for deep-mantle

cycling of these substances. However, he could not test this hypothesis because he did not have data for

oceanic layer 3 and its concentrations of Cl and HzO. SCHILLING et al. (1978) estimated that oceanic

ridge and hotspot volcanism can probably account for the amount of Cl in the surface reservoir. They noted that the present-day outgassing rate is adequate over 4.56 Ga to accumulate a third of the quantity now present in the surface reservoir. They further reasoned that the average integrated transfer rate

throughout the history of the Earth could have been about twice the present rate, and that would account for 65% of Cl in the surface reservoir. They assumed that all the Cl released during arc volcanism origi- nated in the subducted oceanic lithosphere.

Here we estimate the transfer of both Hz0 and Cl from the mantle to the surface reservoir by consid- ering the contributions from subduction of altered

oceanic crust and sediments and from hotspot vol- canism. The transfer of HZ0 and Cl from the mantle to the surface reservoir is commonly thought to occur by igneous activity at the ocean ridges, hotspots, vol- canic arcs, and probably at non-hotspot intraplate volcanoes. The possible return of HZ0 and Cl to the mantle occurs by the subduction of sediments and

altered oceanic crust. For the alteration of oceanic crust, we consider both the seismic layer 2, or the

basaltic layer, and the subjacent layer 3, or the gab- broic and cumulate ultramafic layer.

Our conclusions are summarized in Fig. 1. The ranges of estimates given there reflect various uncer- tainties such as, for example, the role played by the sediment cover on the descending plate (i.e., how much sediment is being carried to what depth). The transfer rates for Cl across the crust-mantle boundary

1613

1614 E. Ito, D. M. Harris and A. T. Anderson. Jr.

14.

12.

; ;10- " 5

6-

. source o sink

WATER 10'4g/yr

FIG. I. Rate terms for geologic processes which act as sources (0) or sinks (0) for annual changes in the amounts of HZ0 and Cl in the crustal reservoir. The ranges shown are based on the estimates given in Table 1. Overlapping ranges for Cl sources and sinks (ordinate) suggest an ap- proximate balance. whereas non-overlapping ranges for HZ0 sources and sinks (abscissa) indicate an imbalance for the geologic processes considered. A. Oceanic ridge mag- matism; B. Hotspot magmatism; C. Arc magmatism; D. Altered oceanic crust; E. Amphibole in altered oceanic crust; and F. Altered oceanic crust and sediments.

appear to be closely balanced using assumed path- ways or geologic processes. However, the transfer rates for HZ0 appear to be in poor balance for the same geologic processes. The mismatch is a factor of

S- 10 and outside the limits of uncertainty. Annually as much as 10 X lOI g of HZ0 may be added to the mantle or returned to the surface reservoir by non- magmatic processes.

Table 1 summarizes the production and subduc- tion rates and average concentrations of Cl and Hz0

Pock3 (Xlo'ig/yr)

+56 +4.4 +0.5 to 10 -56(60.4j' -0.7 to 3.7 -1.8 to 3.5

HZ0 0.2 0.305 2 , to 2 2 2.5 LO 20

(wt.%) C1 48 2905 900 25 to 75 150 430 to 2500

(PPd dL?"SitY 2.9 2.9 2.8 .? . 9 3.1 1.2 to 2.3

k/cm') "Ohm (X1014

20 1.5 1.2 to CU?/YI.)

3.6 0.25 to 1.2 1.5

UliCk"t?SS 6.5 __ 6.5 __ 0.5

(kzn)

used in our estimates. Significant sources of uncer-

tainties affecting our estimates are summarized in

Table 2. Following the order of entry in Table 1, we

first consider the production rate of magmas in dif- ferent tectonic environments. That is followed by an examination of ( I) the average concentrations of Hz0 and Cl in magmas from each environment, and (2) the alteration of oceanic crust. We then look at the

subduction of oceanic crust in order briefly to con- sider both the possible fate of subducted Hz0 and Cl and how different pathways affect the transfer rates

between the mantle and the surface reservoirs. In each section an attempt is made to quantify the un-

certainties and place reasonable upper and lower lim- its on each estimate.

GENERATION RATES OF IGNEOUS ROCKS

The rate of formation of oceanic crust at the ridges was calculated using the figures given in Table I. It is based on the rate of plate creation of 2.995 km2/yr calculated by PARSONS (198 I). The rate we use for the creation of oceanic crust is 56 X lOI g/yr.

SC‘HILLING et al. (1978) have estimated the production rates for hotspot igneous rocks. We use their estimate of 4.35 X IO” g/yr. Eighty percent of the production comes from 6 presently very active, predominantly tholeiitic hot- spots (Iceland, Azores, Reunion-Deccan, Afar, Galapagos and Hawaii). The remaining 20% occurs at 30 currently less active hotspots of more K-rich character. The proportion of oceanic to continental hotspots is not clear. However. from the locations of the hotspots used by SCHILLING et al (I 978) in their estimate, it appears that about one of every three hotspots occurs in continental regions. Oceanic hot- spots are subject to subduction. and their submerged base may be subject to some submarine alteration. The rate we use for the subduction of oceanic crust is thus the sum of 56 X IO” g/yr and K X 4.35 X lOI g/yr, or 58.9 X lOI g/yr.

The estimation of the rate ofcrustal generation by igneous activity in subduction zones is not easy owing to difficulties in quantifying the time averaged rates of extrusion and in- trusion of magmas. The rate of extrusion of volcanic rock. as summarized by SAMPLE and KARIG ( 1982) KARIG and KAY (198 1). and GILL( I98 1). ranges between 2 and 66 km’/ Ma per kilometer of active arc length. The rate of intrusion of batholithic rock was estimated by EVERNDEN and KIS-

used I” Figure 1.

1. Sources Of these est.imates are too “umePO”S to list here, but they are cited in the text. 2. 321% "20 and ,200 ppm Cl were used in Figure 1. 3. Source indicated by +. Sink indicated by -. 4. 60.4~10'~ g,yr Includes ooeanic hotspots. 5. 0.25% "2 and 240 ppm C1 for 6 dominantly thaieiitic hotspota with individual production rate of

0.2 km3/yr. 0.52 H*O and 490 ppm C1 *or 30 dominantly B1LBIiC tlotspots with individual produc- tion rate of 0.01 km3yr.

Cl and Hz0 cycling 1615

TLER (I 970) to be 3.1 km-‘/Ma/km for the Sierras, USA, and by FRANCIS and RUNDLE (1976) to be 2.9 km’/Ma/km for the coastal batholiths of Peru. These estimates give a total rate of crustal generation ranging between 5 and 69 km’/Ma/km. The above estimates also suggest that rates of extrusion are greater or perhaps equal to the rates of intru- sion. However, there is also evidence that arc volcanic suites are complemented by large volumes (3 to 5 times the ex- truded) of mafic to intermediate cumulates (ARCULUS and WILLIS, 1980; STERN, 1979). Independent lines of reason- ing, using volcanic gas emission, leads to the same conclu- sion (ROSE et al., 1982; ANDERSON, in prep.) and energy budgets for explosive volcanic eruptions (HARRIS, in prep.), namely that extruded magma leaves behind about five times its own volume of magma at some depth. The total release of SO2 from arc volcanics ( 1O1’ g/yr; STOIBER and JEPSEN, 1973) divided by average SOz concentration (0.2%, ANDER- SON, 1974) in arc basaltic rock of density 2.8 g/cm’ gives a rate of total volcanism and plutonism of 1.8 X 1 O6 km’/ Ma. Divided by the entire arc length of 3.6 X IO4 km (PAR- SONS, 198 I), this estimate is equivalent to 50 km’/Ma/km, that is to say close to the upper limit of the range derived from the estimates of KARIG and KAY (198 I), GILL (198 I), EVERNDEN and KISTLER ( 1970) and FRANCIS and RUNDLE (1976). Another approach, that of taking the cross sectional area of crust in arcs divided by their probable age, as was done by ANDERSON (1974), gives about 100 km’/Ma/km.

Adopting 2.8 g/cm’ as the average density of arc rock, and 3.6 X lo4 km total arc length (PARSONS, 198 l), the rates of crustal generation derived by all the above methods range from 5 x lOI g/yr to 10 X 1OL5 g/yr, i.e., they vary by a factor of 20. Our preference is for the upper range of 50 to 100 km’/Ma/km which is 5-10 X lOI g/yr (ANDER- SON, in prep.), but subsequent calculations in this paper wilt employ 5 x lOI g/yr as the lower limit and 10 X lOI g/yr as the upper limit. We note that the uncertainty in esti- mating the rate of hotspot generation may be of similar magnitude, i.e., may range from about l/r,, to twice the adopted value.

WATER AND CHLORINE IN MAGMAS

The bulk concentrations of Hz0 and Cl in mid-ocean ridge basalts have been measured for many samples. Vari- ations in quenching pressure, melt composition. proportion of phenocrysts. and proximity to anomalous regions of magma generation (e.g., Iceland. Azores, and Galapagos Islands) may influence the bulk concentrations of Hz0 and Cl (e.g.. MOORE, 1970; HART, 1971; MOORE and SCHILL- ING, 1973; UNNI and SCHILLING, 1978; SCHILLING PI al., 1980). MOORE (1970) estimated, using data on the vesic- ularity of submarine basal& the depth of water (quenching pressure) at which basaltic melts become H20 saturated to be about 500 m (50 bars) for K-poor tholeiites, 800 m (80 bars) for Kilauean tholeiites, and 1800 m ( 180 bars) for alkali-rich basalts from the Revillagigedo Islands. UNNI and SCHILLING (1978) inferred from analyses of Cl in ridge ba- salts that Cl degassing is minimal for samples erupted and quenched at depths greater than 400 m (40 bars pressure). Therefore, most of the Hz0 and Cl in ridge basalts erupted at depths of 2-5 km (200 to 500 bars) should be dissolved within the quenched melt (now glass). The exsolved gas contained in the vesicles in ridge basalts is predominantly CO2 (MOORE et al., 1977). Most of the CO* in submarine tholeiitic basalts is contained within vesicles. with lesser amounts in solution in the glass (see discussion in HARRIS, 1981b). The total amount of CO* can be calculated from MOORE’S (1979) “3-km log-mean vesicularity” data for ridge basalts (70- I200 ppm CO*) and the amount dissolved in the quenched liquid at 3 km (about 180 ppm. see HARRIS, I98 I b). It ranges from 250- 1400 ppm. Since the vesicle gas contains less than 10 mole % Hz0 (MOORE et ul., 1979). it can be inferred that the maximum amount of Hz0 in the vesicles corresponds to less than 0.02% Hz0 in the bulk rock. Therefore. determinations of Hz0 and Cl in some samples of quenched liquids can be used to represent before- effervescence concentrations in the magma. We considered only determinations of Hz0 and Cl in basaltic rocks meeting the following criteria: (I) less than 0.25% KZO, (2) erupted at water depth greater than about 500 m. and (3) less than 10% phenocrysts. The last criterion was necessary for re- ducing the error introduced by correcting for the volume of phenocrysts in arriving at bulk concentrations for H20 and Cl.

The concentrations of Hz0 and Cl in magmas The measured concentrations of HZ0 used in our

probably change before final solidification as a result estimates are summarized in Table 3. Many of the

of processes such as (I) fractional crystallization, (2) measured differences are real and correlate positively

liquid immiscibility, (3) effervescence and formation with an index of differentiation such as K20 concen-

of gas-filled cavities in the upper parts of magma bod- tration (BRYAN and MOORE, 1977). The average con-

ies, (4) addition of Hz0 and Cl from nonmagmatic centration of HZ0 in ocean ridge magmas before ef-

sources, and (5) loss of gas during magma ascent.

Thus, generally, the concentrations of HZ0 and Cl

measured in igneous rocks are different from the con- centrations of undifferentiated magmas. For purposes

of deducing geologic cycles, it is necessary to estimate the bulk concentrations of HZ0 and Cl in magmas. For individual magmas, one must consider the amounts of Hz0 and Cl dissolved in the silicate liq-

uid, the amounts present as an exsolved vapor, and

the amounts structurally bound in phenocrysts. We used petrologic and chemical criteria (e.g., absence of alteration, the modal abundance of phenocrysts,

and concentrations of MgO and K20) for evaluating the significance of rock analyses and to select those

measurements of Hz0 and Cl that were made on the most suitable. least-evolved rocks.

Ocean ridge magmas

1616 E. lto, D. M. Harris and A. T. Anderson, Jr.

Table 3. Concentrations Of “20 in HOAB glasses.

fervescence is estimated to be 0.20 ? .05% (Table 3). At the present time we do not accept the reasoning of DELANEY et al. (1978) that ridge tholeiite melts

contained much less HZ0 (less than .O 1%) at the time of phenocryst growth than at the time of their erup- tion, because HARRIS and ANDERSON (1983a) were

unable to replicate the similar results of MUENOW et al. (1979) for Hawaiian melt inclusions obtained in

the same laboratory as those reported by DELANEY et al. (1978). Finally, we assume that intrusive rocks at ocean ridges (part of layer 2 and all of layer 3) form from magmas identical in bulk chemical com- positions to those that are erupted.

SCHILLING~~~ coworkers (SCHILLING et al., 1978, 1980; UNNI and SCHILLING, 1978; UNNI, 1976) have measured Cl concentrations in nearly 30 glassy basalt samples from the Mid-Atlantic Ridge. The average

of the measured values for samples remote from hot- spots is 48 ppm.

Using the rock production rates we have adopted

(Table I), the contributions of HZ0 and Cl from the generation of oceanic crust at ridges are 1.1 X lOI g HZO/yr and 2.7 X lOI* g Cl/yr. As we mention in

Table 2, there are two major sources of uncertainties in these estimates. First is the possible difference in the concentrations of HZ0 and Cl between glassy and crystalline rock. Second is the accuracy with which we know the average concentrations. We think the concentrations of HZ0 and Cl in crystalline rock would be less than in glassy rock. SCHILLING et al. ( 1980) have 3 samples for which both the glassy rind

and crystalline interior of pillows were analyzed. The concentration of Cl in the crystalline interior is 84 to 89% of the concentration in the glassy rind. It may be premature to generalize from the analyses of 3 samples. At any rate, it is difficult to make use of the generalization since (1) the proportion of glassy to crystalline rock in oceanic crust is not known except that there is more crystalline rock, and (2) it is not certain which concentration is less modified. The average concentration of Cl (and H20) for ridge mag- mas may be higher. Hotspots such as Iceland and the Azores located near or on a ridge contribute elevated concentrations of Cl over some distance @CHILLING et al., 1978, 1980). In addition, there are sections of the MAR such as at 43”N and 45”N where Cl con- centrations are as high as 400 ppm @CHILLING et al., 1980). Such increased concentrations along some sec- tions of ocean ridges may compensate for the un- determined effect of possibly reduced concentration

in crystalline (as opposed to glassy) rock. Overall un- certainty for Hz0 is probably about +25%, and for

Cl about +lOO% and -50%.

Hotspot magmas

SCHILLING et al. (1978, 1980) have measured Cl concen- trations in samples from Iceland, Hawaii, and Azores is- lands. Typical Cl concentrations for the 3 predominant magma types (tholeiites, alkali lavas, and nephelinites) mul- tiplied by their respective magma production rates, give the annual transfer rate of Cl at hotspots of I .29 X lOI g. The measurements on tholeiitic to alkalic glass samples from Hawaii and Loihi by HARRIS ( 1982) give a very similar value of 1.28 X IO’* g (Table I). From the range of concentrations measured by &HILLING et al. (1978, 1980), we estimate the uncertainty in Cl concentration to be about ?50%. Com- bined with the uncertainty in rock production rate, the es- timated range of transfer rate for Cl is from about 0. I3 X IO’* g/yr to about 2.6 X lOI g/yr.

Reliable data are scarce for H20 concentration in hotspot lavas. For example, MOORE (1965) analyzed submarine tho- leiitic lavas from Hawaii and obtained 0.31 to 0.60 wt% H20+. MUENOW et al. (1979) analyzed glass separates from the same rocks studied by MOORE (1965) and found 0.53 to 0.74 wt% H20. HARRIS (1979a, I98 la) reported 0.17 to 0.23 wt% Hz0 in matrix glass of the same or equivalent Kilauean tholeiite samples studied by MOORE (1965) and by MUENOW et al. (1979). The reasons for both previously reported and new discrepancies are not known (c.$, KYSER and JOVOY, 198 I ; RISON and CRAIG, 198 1; MOORE and CLAGUE, 198 1; HARRIS, 1982). However, we feel that the publication of analytical details, instrument “blanks,” sam- ple sizes, and descriptions and valid determinations of an- alytical precision and accuracy, and results for mineral stan- dards will help resolve these uncertainties (e.g., HARRIS, 198la, and unpub. data; KURODA et al., 1982). For sub- marine lavas, MOORE (1970) found a direct relationship among HzO+, Cl, F, P205. and K20 in fresh tholeiitic to alkalic pillow lavas. Similarly, co-variation among H20, Cl, PzOs, and K20 in tholeiitic to alkalic glass samples included in olivine or as quenched rims of pillow basalts were re- ported by HARRIS ( 1982). We assumed that the relationships found by HARRIS (1982) for Kilauea and Loihi are repre- sentative of undegassed glassy lavas from other hotspots.

We use the relationships found by HARRIS ( 1982), and the proportion of predominantly tholeiitic (80%) to predominantly alkalic (20%) hotspot igneous ac- tivity (SCHILLING et al., 1978), to make a preliminary estimate of HZ0 in hotspot magmas: predominantly

tholeiitic magmas contain about 0.25 wt% Hz0 and predominantly alkalic magmas contain about 0.50 wt% HZ0 before effervescence. Consequently, Hz0 transferred from the mantle to crust at hotspots is 1.3 X lOI g/yr. Our undoubtedly subjective estimate of uncertainty is similar to that for Cl, i.e., about a factor of 2 either way. The use of the relationships found by MOORE (1970; 0.40 wt% H20 for tholeiitic; 1 .O wt% for alkalic) leads to a 75% increase in the esti- mated transfer rate. The increase is within the un- certainty.

Subduction zone magmas

Our aim here is to estimate the average HZ0 and Cl concentrations of typical subduction zone mag- mas prior to differentiation. The problems surround-

Cl and HZ0 cycling 1617

ing the genesis of andesitic magmas severely com-

plicate this task. Several lines of evidence suggest that andesites are derived from subduction zone basaltic

magmas (e.g., KUNO, 1968; STERN, 1979; ANDER- SON, 1979, 1982; ARCULUS and WILLIS, 1980). Fur- thermore, experimental evidence indicates that an-

desite cannot be derived as a primary magma from subducted oceanic crust (wet or dry) at a depth of 100 km and that andesite is unlikely to be a primary magma generated by the melting of wet upper mantle

peridotite at about 30 kb (e.g., WYLLIE, 1978). Moreover, talc-alkaline magmas that become sat-

urated with an HzO-rich fluid phase are likely to crys-

tallize and form intrusions at shallow levels within the crust (SCARE et af., 1981) and should not be observed as lavas. Hence, on the basis of our own studies and those of others, we conclude that andes- itic magma is produced by differentiation of basaltic magma. Accordingly, our estimates of HZ0 and Cl concentrations in subduction-zone magmas are based primarily on studies of basal& although the concen-

trations of Hz0 and Cl in liquids of andesitic com- position were also considered.

EGGLER ( 1972) estimated from his experiments that an- desitic magma from Paricutin, Mexico, originally contained 2.2 f 0.5 wt% H20. SEKINE et al. (1979) deduced from their experiments on rocks from Asama and Sakurajima that a melt corresponding to the groundmass composition (69 wt% SiOz) contained 3.3 wt% HZ0 just before an early explosive phase of one eruptive sequence, whereas the groundmass liquids of later lava flows contained less Hz0 (as little as 0.75 wt%). Water dissolved in silicate melt inclusions in phenocrysts have been estimated for some basal& basaltic andesites. and andesites from circum-Pacific volcanoes (e.g , ROSE et (11.. 1978; ANDERSON, 1979, 1981). The estimated Hz0 contents generally range from 2 to 4 wt%. Develop- ment of a new vacuum fusion technique (HARRIS, 198 la) permitted direct determinations of Hz0 concentration in melt inclusions in individual crystals (HARRIS, 1979, 198 lc, and unpub. data; HARRIS and ANDERSON, 1983b) from Fuego (1974 eruption), Pagan (southeast rift of Mt. Pagan), and Agrigan (Kimi parasitic cone). Since the same melt inclusions have also been analyzed by electron microprobe for major elements, S, and Cl, the actual glass compositions are known. The measured Hz0 contents are about the same for both hypersthenic and pigeonitic melt inclusions (about I .3 to 3.4 wt%). The concentrations of Hz0 in the magmas of compositions identical to the melt inclusions are not known. However, the common occurrence of hornblende in olivine-bearing gabbroic rocks seems consistent with 2 to 4 wt% of Hz0 in magmas (ANDERSON, 1980). Based on these studies, we assume that plutonic and volcanic rocks in subduction zones form from parental basaltic liquids with similar Hz0 contents and estimate that 2.0 f 1 .O wt% is a representative bulk Hz0 value for most medium-KzO sub- duction-zone basaltic liquids before differentiation and ef- fervescence.

The average concentrations of Cl in subduction zone magmas have been estimated from the concen- trations of Cl in silicate melt inclusions in olivine phenocrysts (ANDERSON, 1974, 1982; ROSE et al., 1978; HARRIS, 1979b; HARRIS, 198 lc). ANDERSON (1982) used a relationship between Cl and K20 con- centrations to identify the melt inclusions that best represent the parental magma for each of some 8

subduction zone volcanoes. The Cl concentrations

range from 400 to 1400 ppm with an average of 900

ppm (ANDERSON, 1979, 1982). The total annual contribution of Hz0 and Cl from

igneous activity at subduction zones is 1 .O$:& X lOI g Hz0 and 4.5’2:: X 10” g Cl. The very large un-

certainty is due in most part to the wide range of estimates of the crustal generation rate at arcs.

WATER AND CHLORINE IN ALTERED CRUST

ITO and ANDERSON (1983) and ITO and CLAYTON

(1983) have studied the alteration of gabbroic rocks from the Mid-Cayman Rise in the Caribbean. The

rocks they studied were collected directly from the rift valley wall by the submersible ALVIN. The knowledge of precise location of sampling enabled ITO and ANDERSON (1983) and ITO and CLAYTON (1983) to infer that in slow-spreading ridges (1) al- teration was intimately related to deformation, (2)

significant seawater penetration was limited to about upper 1 km of layer 3 rocks, and (3) the abundance of hydrous minerals decreases with the depth into oceanic crust. This section of the paper relies heavily on the results of their studies in discussing the alter- ation of slow-spreading ridges. In addition, we have relied on studies by MOTTL (1983) PRICHARD and CANN (1982) HONNOREZ (1981) HUMPHRIS and THOMPSON ( 1978), HELMSTAEDT and ALLEN (1977) BONATTI (1976) BONATTI et al. (1971, 1975) HODGES and PAPIKE (1976) MUEHLENBACHS and CLAYTON (1972a,b), and AUMENTO and LQUBAT (197 1). Studies by ANDERSON et al. ( 1983) EDMOND

et al. (1979, 1982) and STAKES and O’NEIL (1982) in particular, were used in the discussion of alteration

in fast-spreading ridges. We have also made use of results of studies on ophiolites, most notably the stud- ies by GREGORY and TAYLOR ( 198 I), COCKER et al. ( 1982), and STERN and ELTHON ( 1979). In general,

we have placed greater reliance on those studies

which had good stratigraphic control on sample lo- cation within vertical sections of crust.

Throughout this section and the rest of the paper, “oceanic crust” refers to the basement or seismic lay- ers 2 and 3 and excludes the sediment cover.

Hydration of oceanic crust

The oceanic crust gains HZ0 by submarine alter- ation. The HZ0 gain is controlled by the degree of alteration and the resulting mineralogy. Chlorite, ser- pentine, smectite, illite, and amphibole are the prin- cipal hydrous minerals in the altered oceanic crust. There are several ways of estimating the amounts of HZ0 fixed in the oceanic crust. The most direct, but perhaps least constrained, approach is to analyze the bulk H20 of rocks. This method provides no infor- mation regarding the storage sites for HZ0 within the rock. Another approach involves the abundance and species of hydrous minerals. The concentrations of

1618 1E. Ito, D. M. Harris and A. T. Anderson, Jr.

elements such as CL (and F) which could substitute for OH in minerals is generally very low (I I50 ppm, ITO and ANDERSON, 1983). Although the error in- troduced by assuming stoichiometric J&O content in h~drD~SS%caQ?SknOt necesW%y SXnti (BOETTCHER and O’NEIL, 1980), the likelihood of finding sub- stantial 0 in OH site should be small for hydrous minerals which were formed in submarine environ- ment. The third method is to use the direct relation- ship that exist between the increase in oxygen isotopic composition and increase in F&O* of whole rocks which were altered by halm~olysis (T I 150°C; e.g., MUEHLENBACW and CLAYTON, 1972a). This last method, however, works only for the rocks altered by halmyrolysis (T I 15O’C) and not for the rocks altered hydrothermally. All these three methods, however, depend on knowing where the sample came from, i.e., the depth below the basement floor. A method which does not require this information is the mass balance approach which uses the compo- sition of hydrothermal vent fluids (e.g., EDMOND et al., 1979, 1982). When combined with the results of experiments on rock-water interactions (e.g., SEY- FRIED and Mont, 1982; SEYFRIED and BISCHOW, t 98 L ; Morr~ and SEYFRIED, 1980; MOTTL and HOLLAND, 1978; MOTTL d al., 1979) and the ob- served mineralogy of natural samples, this method allows one to arrive at some estimate of the amount of secondary minerals formed in the oceanic crust.

We make use of the second and the fourth approaches. Up to now, the mineralogy ofthe altered oceanic rocks has been studied more extensively for samples from slow- spreading ridges such as the Mid-Atlantic Ridge, the Mid- Cayman Rise, and the Carlsberg Ridge (e.g.. HUMPHRIS and THOMPSON, 1978; HODGES and PAPIKE, 1976; AUMENT~ and LOUBAT, 1971; ITO and ANDERSON, 1983; HELMS- TEADT and ALLEN, 1377; CANN, 1969, 1971; BONATTI, 1976; BONATTI et al., 197 L, 1975; MIYASHIR~ er al.. 1971, 1979). The hydrothermal vent fluids, on the other hand, have only been discovered in fast-spreading ridges, the Ca- lapagos Rise and the East Pacific Rise (EDMONDet al., 1979, 1982). Hence, there is some ambiguity in combining the mineralogic data from the slow-spreading ridges with the

solution data from the fast-spreading ridges. Nevertheless, this may be the best approach available to us at present (SEYFRIED, in prep.). Our estimate of the mineralogy and the amount of hydration in the oceanic crust is summarized in Table 4 and in Fig. 2. As can be seen from Table 4, the bulk Hz0 concentration of altered oceanic crust is about 1 S f 0.5 wtb, The abundance of hydrous minerals is more or less consistent with the results of the studies cited above including the finding by ITO and ANDERSON (1983) of a decrease in the pervasiveness of alteration at greater depths. Generally, dredged rocks are extensively altered, probably because of biased sampling toward fault scarps and fracture zones. In comparison, cored rocks (HODGES and PAPIKE, 1976; HELMSTAEDT and ALLEN, 1977; ANDERSON el ai., 1983) and rocks from the Mid-Cayman Rise are not so extensively altered. The solution data (EDMOND& al., 1979, 1982; SEYFRIED, in prep.}, on the other hand, require an extensive alteration of oceanic crust, one equivalent to the alteration of the entire layer 2 into greenstones. Thus our estimate represents a compromise between the “require- ments” indicated by different types of data.

Our estimate of 1.5 i 0.5 wt% Hz0 for the altered oceanic crust gives, when multiplied by the annual

rate of rock subduction (Table l), 8.8 5 2.9 X 1OL4 g/yr as the total amount of Hz0 subducted each year. This is approximately 9 times the amount of f1*0 contained in unaltered oceanic crust. We did not in- clude in our consideration possible additional sites of Hz0 storage such as fluid inclusions and Hz0 con- tained in intergranular spaces. The inclusion of these additional storage sites is not presently possible in any quantitative way (Table 2). However, the addi- tional Hz0 in these sites might significantly increase our estimated figure.

Cldorine in oceanic crust

In comparison to HzO, the gain or loss of Cl in the oceanic crust with submarine alteration is difficult to estimate. The di@iculty arises mainly from the pau- city of data. Published analyses of Cl in altered oceanic rocks are, at present, Iimited to those of gab- bros from the Mid-Atlantic Ridge (HONNOREZ and KIRST, 1975; HODGES and PAPIKE, 1976; PRICHARD

Table 4. Estimated abundance and H20, Cl concmtrations of various rock types in altered oceanif crust?

Thickness -Em* Hydrous Average2 Average' Bulk3 Bulk' (km) TYPO MiWW~l8 H20 (I) Cl (ppm) H20 (%I Cl (ppm)

0.5 + 0.3 Altered by 30% Clay 20 $6 6.0 15 Halmyrolysis

1.0 + 0.5 Greenstone $09. Chlorite 12 ii 5.4 42.5 15% Amphibole 2 100 15% Epidate ,? i:!

0.5 C 0.3 Gnphibolitc 50% Amphibole P 150 I 75

7 Serpentinite 100% Serpentine 12 >O ? ?

6.5 Total crust 1.59.5 5095

lThis table forms the basis for the entries far the altered MORB in Table 1. 2Average H$ and Cl mnfentratiom of each hydrom mineral. It vas assumed that anhydrous minerals contain no H$ and Cl. Furthermore, H$J and Cl oontained in fluid inclusions and in intergranular spaces are neglected. The concentratian of CL in the hydrous minerals far which Cl was not detected by electron micromwbe was asswsd to be one-half the detentlan level tsee text) and is shown-in italics. 3&O and Cl concentrations for each Indicated rock type. 4The uncertainties given are the sums of rather subJective estimates discussed in differant Darts of the text. The 5ourcez of uncertainties listed in Table 2 apply here as well.

Cl and Hz0 cycling 1619

and CANN, 1982) and from the Mid-Cayman Rise (ITO and ANDERSON, 1983). This rather limited data base suggests that most minerals in submarine gab- bros do not contain detectable concentrations (> 100 ppm) of Cl with the exception of amphibole and apatite. Under the routine analytical conditions, Cl concentrations of less than 100 ppm are not reliably measured by an electron microprobe using wave- length dispersive spectrometer. Hence, hydrous min- erals such as chlorite, mica, and talc, or perhaps even anhydrous minerals, may in fact contain some Cl. For our purposes, we can assume either (1) on the average ail minerals for which Cl was not detected by microprobe contain 50 ppm Cl, or (2) on the av- erage only hydrous minerals for which Cl was not

detected contain 50 ppm Cl, or (3) on the average Cl concentrations in all minerals except amphibole and apatite are negligible. Case (2) has been assumed in constructing Table 4.

Additional sites of Cl storage in the altered oceanic crust are fluid inclusions and intergranular spaces. Fluid inclusion in hydrothermal quartz typically con- tain very saline solution (JEHL et al., 1977), so that the amount of Cl is probably not negligible. However, for the present we cannot realistically assess this quantity, and it has not been included in Table 4.

The concentrations of Cl in amphibole and apatite are variable. In addition, the modal abundance of these two minerals are variable from rock to rock. For example, apa- tite is rare in the gabbros from the Mid-Cayman Rise (Ho and ANDERSON, 1983) and would only cont~bute a max- imum of 30 ppm Cl toward the bulk rock Cl concentration. On the other hand, apatite is common in the gabbros from the Equatorial Mid-Atlantic Ridge and could perhaps con- tribute substantial amounts of Cl (HONNOREZ, pers. com- mun.). Similarly, amphibole is abundant in some altered rocks, and its Cl concentration ranges from about 100 to 10,000 ppm or more (HONNOREZ and KIRST, 1975; PRI- CHARD andCANN, 1982;Iroand ANDERSON, 1983; HON- NOREZ et al.. 1983). Among the Cl-poor amphiboles, horn- blende and actinolite contain an average of 150 ppm and 100 ppm Cl respectively (ITO and ANDERSON, 1983). Cl- rich amphiboles typically occur either in veins and sur- rounding vugs in metagabbros or in amphibolitized gabbros which have not acquired a gneissic texture (ITO and AN- DERSON, 1983; HONNOREZ et al., 1983). In rocks from the Mid-Cayman Rise. the CT-rich amphiboles occur in a 100

WATER wt %

c CHLORITE I

UNALTERED CRUSl

FIG. 2. Amount of various hydrous minerals in altered oceanic crust and average concentration of H20 in each mineral.

CHLORINE mm

50 100 150 0

SMECTITES AND ILLITES

l- ~CHL~RITE I

I I

I I

I I

FIG. 3. Amount of various hydrous minerals in altered oceanic crust and average concentration of Cl in each min- eral.

m interval at the transition between rocks with abundant amphibole and rocks with little amphibole (ITO and AN- DERSON, 1983). The abundance and the distribution of Cl- rich amphiboles in other ocean ridges are not known, and hence the contribution of Cl-rich amphiboles has been ig- nored. We have, however, assumed that amphiboles in the amphibolite layer are all ho~blende and contain 150 ppm Ci (Table 4).

As can be seen from Table 4 and Fig. 3, the bulk Cl concentration of altered oceanic crust is 50 I 25 ppm. The estimate for Cl is more indirect than that for Hz0 because of its dependence on the estimated abundance of various hydrous minerais in the altered oceanic crust, and it is therefore less reliable. Our estimate of 50 + 25 ppm Cl for the altered oceanic crust gives, when multiplied by the annual rate of rock subduction (Table I), 2.9 ? 1.5 X lOI* g/yr as the total amount of Cl subducted each year. This figure, which does not include possible cont~butions from fluid inclusions and intergranular spaces, is es- sentially the same as that produced by oceanic ridge magmatism.

SUBDUCTION OF OCEANIC CRUST AND POSSIBLE GEOLOGIC CYCLING OF

WATER AND CHLORINE

The main items we consider here are (1) what is meant by subduction, (2) the role of sediments, and (3) the stability of minerals in the subduction zone environment. We assume that the subduction rate of oceanic crust balances the emplacement rate of new crust at oceanic ridges and oceanic hotspots (Table 1).

The word “subduction” as generally used by ma- rine geologists apparently means “the disappearance of material beneath the bottom of the sea at the trenches.” In discussing possible cycling of Hz0 and Cl, a distinction must be made between subduction into crustal depths or subduction into mantle depths. Available geophysical techniques do not permit this distinction and hence we attempt to address both possibilities, especially in the case of sediments.

1620 E. Ito, D. M. Harris and A. T. Anderson, Jr

Subduction of sediments

KINSMAN (1975) has proposed that rifted or trail-

In some subduction zones, for example in most of the western Pacific, the sediment layer (100 to 500

ing continental margins accumulate a very thick

m thick) seems to be subducted in the sense that it disappears beneath the trenches (SCHOLL et al., 1977;

prism of sediments and eventually become subduc-

HAWKINS et al.. 1979; LEGGETT, 1980). There is some evidence that some components of sediment

tion zones. Trailing continental margins were sites

are subducted deep enough to be refluxed at arc vol- canoes (BROWN et al., 1982). Well-compacted, fine-

of initial rifting and in some places have accumulated

grained oceanic sediments under kilometers of over- burden may still have 30% porosity or greater (e.g.,

thick beds of evaporites (KINSMAN, 1975; BURKE,

BERNER, 1973). Therefore, subduction of a 500 m thick layer of sediments with 30% porosity at all sub-

1975; BONATTI et al., 1970). If some subduction

duction zones, if the pore water has normal seawater composition, can carry to some unknown depths in

zones have formed as proposed by KINSMAN ( 1975),

the subduction zone amounts of Cl and HZ0 equal to those produced by all the igneous activity each

large quantities of Cl may have been subducted; but

year (Table 1). If the pore water is not carried beyond the accretionary prism and only structurally bound

his proposal remains unproven. At present, the only

(and adsorbed) Hz0 and Cl are subducted deeper (CARSON, 1977; ARTHUR et al., 1980), these amounts

area where subduction of evaporites is plausible is in

are greatly reduced. If the weighted Cl and HZ0 con- centrations in clays and carbonates from geosynclinal

the Hellenic Trench in the Mediterranean Sea. How-

units (6.3 wt% HZ0 and 1200 ppm Cl; RONOV and

YAROSHEVSKIY, 1976) are assumed to represent Cl

ever, Messinian evaporites are known to accumulate

and Hz0 in subducted sediments, the total Cl and Hz0 associated with the subduction of sediments are

in the Hellenic Trench and are not subducted (LEG-

3.2 f 1.0 X 10” g and 1.7 f 0.5 X lOI g each year respectively, depending on the sediment thickness

GETT, 1980).

and density (Fig. 1 and Table 1).

During the subduction of oceanic crust and the overlying sediment the main changes that take place down to a depth of about 15 to 18 km are dewatering and diagenesis of the sediments. Expelled water may percolate along the faults and fractures in the accre- tion wedge and return to surface. Clay minerals, if sediment is carried down deeper, will dehydrate at 20 to 50 km depth depending on the temperature distribution (DELANEY and HELGESON, 1978; ERNST, 1976, 1977) and the bulk composition of the de- scending crust. The blueschist-facies and granulite- facies metasedimentary rocks which are intercalated with metabasalts of similar metamorphic grades in fossil convergent-plate margins indicate that in some

subduction zones sediments were carried down to

depths of 30 to 70 km (ERNST, 1976, 1977; EVANS

and TROMMSDORFF, 1978; HOLLAND, 1979; EN- GLAND and HOLLAND, 1979; MARUYAMA and YA-

MASAKI, 1978; BLACK and BROTHERS, 1977). This evidence requires that in some subduction zones sed-

iments are carried to depths greater than 45 km. The amounts of Hz0 and Cl contained in metamor-

phosed sediments at such depths are probably similar to or less than those found in blueschist rocks (2 to 4 wt% H20; ERNST, 1963).

Subduction and geologic cycling qf HI0 and Cl

The stabilities of minerals, however, are deduced from combinations of different experimentally de- termined phase equilibria for Fe-free systems. Min-

erals actually present in the subducted slab, because they contain Fe, will break down at shallower depths or lower temperatures than indicated by phase equi-

Among the hydrous minerals still present in the

subducted oceanic crust and sediments at mantle

libria studies of Fe-free systems. Furthermore, large

depths, those other than micas will dehydrate in the

uncertainties are introduced when data from 3- or

interval 60 to 100 km (DELANEY and HELGESON, 1978). Most of the chlorite which formed during sub-

4-component systems are combined to deduce the

marine metamorphism probably did so at tempera- tures below 500°C and is likely to be metastable (HELGESON et al., 1978). Consequently, these will

stability fields of minerals in more complex systems.

break down at shallower depths than either hom- blende (- 100 km and

Such uncertainties and unknown rates of migration

-95O”C, e.g., AL.LEN and

of HZ0 and other volatiles in relation to rates of sub-

BOETTCHER, 1983) or stable forms of chlorite such as clinochlore (- 110 km and -850°C. STAUDIGEL

duction indicate that we have only a crude idea of

and SCHREYER, 1977; JENKINS, 1981). Mica is more likely to melt at its high temperature limit (- 11 OO”C, WY LLIE, 1977) than to dehydrate at its high pressure

(the sequence and extent) of the mineralogic reac-

limit (- 150 km, WYLLIE, 1977) depending on the

tions that occur in the subducted crust.

Y-7’ trajectory of different subduction zones.

The depth to the top of the high Q, inclined slab beneath active volcanoes is about 120 to 150 km in most volcanic arcs (e.g., KATSUMATA and SYKES, 1969; JAMES, 197 1; ENGDAHL, 1977; ISACKS and BARAZANGI, 1977), or about 50 to 100 km below the depth where most hydrous minerals originally pres- ent in the altered oceanic crust decompose to vapor and anhydrous minerals. There is probably not enough mica in the slabs to supply the Hz0 contained in arc magmas. Mica also contains little Cl (SMITH et a/., 198 1). However, if Hz0 expelled from dehy- drating minerals rises sufficiently slowly through the subducted crust and the overlying mantle, or if ex- pelled Hz0 reacts with the anhydrous minerals either

Cl and Hz0 cycling 1621

in the slab itself or in the overlying mantle, then Hz0 can be carried down to greater depths than predicted solely from mineral stabilities (e.g., by induced back- arc convection, MCKENZIE, 1969; SLEEP and TOK- SOZ, 197 1). Consequently, the estimated stabilities of the original hydrous minerals during su~uction need not be the only important considerations if the released Hz0 stays within or near the slab during subduction to depths of at least 100 km.

Assuming that during subduction H20 and Cl stay within the altered oceanic crust down into the man- tle: the sum of the sources (i.e., magmatic activity at ocean ridges, hotspots, arcs) and sinks (i.e., subduc- tion) appears to be negative for Hz0 due to subduc- tion and positive for Cl due to net production (Table 5). If sediments are not subducted into the mantle, as calculated in the previous section, 8.8 * 2.9 X lOI g of HZ0 and 2.9 rt 1.5 X lOI g of Cl are subducted annually. For H20, the e,rcess of subduction over production is 6.6?::: X lOI g/yr, or 1.2-I 1 X lOi g/yr (i.e., l- 11 times the rate of addition of HZ0 to the crust through oceanic ridge magmatism). For Cl, it is about 40% of the produced amount. If sediments are subducted into the mantle (500 m thick; 2-4 wt% HzO, ERNST, 1963; maximum of 1200 ppm Cl, RONOV and YAROSHEVSKIY, 1976, Table 3, mean geosynclinal sediments; density 1.2-2.3 g/cm3), then the total amounts subducted become: 9.7 -t 3.4 X lOI g of HZ0 and 6.1 & 2.5 X 1Or2 g of Cl per year. For HzO, the difference between sediment-no sediment is small. However, for Cl the difference could be quite large. The concentration of Cl we have used leads to the doubling of the total Cl subducted bringing the total subduction to three-quarters (3/4) of the produced amount.

Compared to the amounts of H20 and Cl produced by arc magmatism, the amounts subducted are about 9 times (range 2-230) larger for Hz0 and about the same (within rt 50%) for Cl (Table 5). The balance for Cl is consistent with refluxing of subducted Cl by

arc magmatism (SCHILLING et d., 1978) but does not require it. The imbalance for subduction zone re- fluxing of Hz0 contained in altered oceanic crust is not due to a series of errors made in our estimation. While we grant that the imbalance may be smaller than our estimate (i.e.. the production rate of HZ0 by arc magmatism may be higher and the amount of H20 subducted into the mantle may be lower) it is nevertheless likely to be real. The idea that seawater is disappearing into the mantle is not new (e.g., DICK- INsoNand i.AJTH, 197 1; FYFE, 1978). Our result lends support to their hypothesis. We estimate the net rate of disappearance to be 1.2-11 X lOi g/yr, or 0.75- 8.3% of the new crustal inventory per 100 Ma. In as much as these rates would suggest major variations of sea level, we feel that the higher rates in this range are geologically unrealistic. However, net subduction rates of l-3 X iOr4 g of HZ0 per year for periods of 100 Ma or more may be consistent with the geologic record.

SUMMARY

We have considered the effects of the alteration of oceanic crust and the subduction of oceanic sedi- ments on the cycling of HZ0 and Cl between the mantle and surface reservoirs. The effect of the al- teration of oceanic crust on Hz0 cycling is significant, but it is negligible for Cl cycling. The formation of hydrous minerals in oceanic crust during the sub- marine alteration leads to approximately an order of magnitude increase in the concentration of H20. The subduction of sediments into the mantle has little effect on HZ0 cycling (average sediment thickness 500 m) largely because the effect of the alteration is much larger. In contrast, the subduction of sediments into the mantle has a relatively large effect on Cl cycling. This is because the total amount of Cl con- tained in sediments is about the same as that in al- tered oceanic crust, i.e., the subduction of sediments will double the amount of Cl being subducted.

Our conclusion, summarized in Fig. 1 and Table 5, reflects the sum of various uncertainties (Table 2). On the basis of estimated rates and concentmtions derived from a reasonable set of assumptions, we conclude that for Cl production and subduction are in balance or there is net production, whereas for HZ0 there appears to be net subduction. The reflux of subducted HZ0 to the surface reservoirs is not fully accounted for by (I) igneous activity at volcanic arcs in subduction zones, (2) mantle Hz0 associated with igneous activity at oceanic spreading centers and hot- spots, or a combination of (1) and (2).

,4cknow[~~g~m~ffts-An early version of the paper formed a chapter of lto’s Ph.D. thesis. A similar, early version was presented by Harris at 1979 IUGG Assembly held at Can- berra, Australia. Two later versions were presented by Ito at the Seventh Symposium on Geochemical Cycles held in 1980 at Atlanta, GA, and at the LPI Topical Conference on Planetary Volatiles held in 1982 at Alexandria, MN. Yet another version was presented by Anderson at the 1982

1622 E. Ito, D. M. Harris and A. T. Anderson, Jr,

AGU Spring Meeting. Parts of the paper were prepared dur- ing Ito’s stay as a Carnegie Fellow at the Department of Terrestrial Magnetism, Carnegie Institution of Washington, and were finally completed at the University of Minnesota. Harris gratefully acknowledges support from American Geophysical Union and International Association of Me- teorology and Atmospheric Physics (IAMAP) used to defray travel expenses to attend the IUGC meeting in Canberra. The work was partially supported by NSF grant, No. EAR 76-15016 to Anderson (University of Chicago). We thank J. G. Schilling, J. R. Delaney, J. Honnorez, and an anon- ymous reviewer for constructive comments.

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