3.15 hydrothermal alteration processes in the oceanic …lrg.elte.hu/oktatas/szubdukcio es kopenyek...

3.15Hydrothermal Alteration Processesin the Oceanic CrustH. Staudigel

University of California at San Diego, La Jolla, CA, USA

3.15.1 INTRODUCTION 511

3.15.2 THE UNALTERED OCEANIC-CRUST PROTOLITH 5123.15.2.1 A “Standard Section” for the Oceanic Crust 5123.15.2.2 Estimating Unaltered Oceanic Crust Compositions 515

3.15.3 DETERMINING THE ALTERED COMPOSITION OF THE OCEANIC CRUST 5173.15.3.1 Recovery Rate 5173.15.3.2 Types of Alteration 5183.15.3.3 Duration of Alteration 5193.15.3.4 Determining the Composition of Extremely Heterogeneous Altered Crust 520

3.15.4 CHEMICAL CHANGES IN ALTERED CRUST COMPOSITION DUE TO HYDROTHERMALPROCESSES 522

3.15.4.1 Time Dependence of Crust Hydration and Carbonate Addition 5223.15.4.2 Chemical Fluxes between Oceanic Crust and Seawater: Methods and Uncertainties 5233.15.4.3 Chemical Fluxes 524

3.15.5 DISCUSSION 5293.15.5.1 Hydrothermal Fluxes: Rock Data versus Fluid Data 529

3.15.5.1.1 Uncertainties 5303.15.5.1.2 Bulk fluxes 5303.15.5.1.3 Reconciling hydrothermal fluxes from fluid and rock data 531

3.15.5.2 Impact of Ocean-crust Composition on Arc Processes and Mantle Heterogeneity 531

3.15.6 CONCLUSIONS 532

REFERENCES 533

3.15.1 INTRODUCTION

Hydrothermal alteration processes occurring inoceanic crust impact the physical, chemical, andbiological processes of the Earth system. Thesehydrothermal systems are manifested in ventsranging from 350 8C black smokers, foundexclusively in the axial zone of some ridgesegments, to 20 8C low-temperature vents at theridge axis or flanks. Collectively, these systemsare responsible for ,20% of Earth’s total heatloss (11 TW; C. A. Stein and S. Stein (1994a,b))and have major impact on ocean and solidearth chemistry. Elderfield and Schultz (1996)estimate black-smoker water fluxes to be,3.5 £ 1012 kg yr21 and low-temperature fluxesto be ,6.4 £ 1014 kg yr21 (at 20 8C). These

hydrothermal fluxes also carry substantialelemental flux between seawater and the oceaniccrust. Combined with ocean-crust generation andrecycling, these processes produce a two-waygeochemical pathway between the oceans and themantle. Recycling of altered oceanic crust into themantle is likely to produce some of the mantle’schemical heterogeneity (e.g., Hofmann, 1988; seeChapter 2.04) and the delivery of mantle-derivedmaterials to seawater through hydrothermalsystems has profound effects on seawater chem-istry (e.g., Wheat and Mottl, 2000; Chapters 3.15and 6.07). Hydrothermal vents in mid-oceanridges offer a unique habitat for very diversebiological communities that derive much of theirenergy needs from chemical energy in vent fluids(Jannasch and Mottl, 1985; Jannasch, 1995).

511

The interior of the oceanic crust is likely to host adeep-ocean biosphere that reaches to at least500 m depth (Furnes and Staudigel, 1999).

It is important to quantify hydrothermal chemi-cal fluxes because they bear on the chemical andbiological evolution of the Earth, the chemicalcomposition of seawater, geochemical massbalance at arcs, and the heterogeneity of themantle. Hydrothermal fluxes can be independentlydetermined by analyzing the composition ofhydrothermal fluids or by analyzing the alteration-related chemical changes in the oceanic crust.Ideally these two methods should yield the sameresults, but a comparison of data shows that thereare major discrepancies between these types ofestimates (e.g., Hart and Staudigel, 1982; Chapter3.15). Reconciling these discrepancies is impor-tant for improving our understanding of thiscentral theme in Earth system sciences.

This review focuses on chemical flux estimatesderived from studies of the oceanic crust, explor-ing in detail how such estimates are made, and theunderlying assumptions and uncertainties. Threemain themes will be covered. The first focuses therole of the original igneous characteristics of thecrust in determining the nature of hydrothermalalteration processes. This includes how primarylithology and composition influence alteration,and difficulties encountered in determining anunaltered “fresh-rock” baseline composition forany particular ocean-crust section. The secondtheme focuses on the methods by which the bulk-altered oceanic composition is determined, andthe attendant uncertainties. These include thedifficulty of determining an average compositionof a very heterogeneous medium by the analysesof rather small samples, and the limitationsimposed by an incomplete sampling process onthe ocean floor. Finally, hydrothermal fluxesinferred from ocean-crust data are compared tofluxes from hydrothermal vent studies and thereasons behind their differences are explored.

3.15.2 THE UNALTERED OCEANIC-CRUSTPROTOLITH

Determining the original igneous characteris-tics of altered oceanic crust is an important step inunderstanding its alteration. There are two verydifferent reasons for this.

First, the primary igneous characteristics of theoceanic crust determine important aspects ofalteration by controlling heat exchange, fluidflow, and the reactivity of crustal materials withhydrous fluids. These primary characteristics mayvary as a function of spreading rate, ridgemorphology, or position in a volcano, producinga range of alteration types. Effectively integratingthese different crust types is an important

challenge in producing a meaningful globalintegration of alteration fluxes. This problem isaddressed here by defining a “standard” sectionfor normal oceanic crust based on results fromavailable sample materials.

Second, the original composition of oceaniccrust plays an important role in determiningchemical fluxes. Uncertainties in the originalcomposition translate directly into uncertaintiesof chemical fluxes, and in many cases thesedifferences are the main sources of errors. Both ofthese features are addressed in the followingsections.

3.15.2.1 A “Standard Section” for the OceanicCrust

The most comprehensive understanding ofprocesses occurring in the oceanic crust comesfrom the study of subaerially exposed ocean-crustsections in ophiolites, which offer extensive andoften continuous exposure over very large areas.The observations from ophiolites complement thesignificantly more limited data from submerged,in situ oceanic crust, which in turn serves as animportant test for the validity of these subaerialanalogues. For this reason, ophiolites are key tounderstanding the relationships between alterationbehavior and primary crust characteristics. Themain ophiolites and ocean-crust drill holesdiscussed in this paper include the Troodosophiolite in Cyprus, the Samail ophiolite inOman, and ocean-crust drilling sites 332B and417/418 in the Atlantic Ocean and sites 504Band 801C in the eastern and western Pacific,respectively (Figure 1, Table 1).

The classic view of the oceanic crust is based onthe “Penrose” ophiolite assemblage (PenroseConference Participants, 1972) that includes acharacteristic sequence of rock types: basal maficcumulates are overlain by gabbro-norites, gab-bros, sheeted dikes, and pillow lavas on the top.Pillow lavas may be overlain by hydrothermalmetalliferous sediments, that are commonly over-lain by pelagic, typically siliceous fine-grainedsediments (cherts). This three-decade-old ophio-lite model of the oceanic crust remains effectivelyunchanged in terms of the main rock types foundin oceanic crust and in terms of their stratigraphicsuccession. The main modifications to this modelinclude changes in the average thicknesses ofunits. In addition, there is some in situ oceaniccrust that appears to lack any extrusives, but thistype of ocean crust is mostly confined to very slowspreading ridges and represents probably ,5% ofthe total crust produced (e.g., Bach et al., 2001).

Study of the global impact of ocean-crustgeneration and alteration ideally requires inte-gration of data from all major crustal types, ifpossible, including several studies that offer

Hydrothermal Alteration Processes in the Oceanic Crust512

independent constraints. Unfortunately, there aretoo few sections studied for such coverage, and forthis reason one has to combine all observations intoone “general” type of ocean-crust section. This isclearly an over-simplification, but to a first order agood approximation for the most abundantlyproduced oceanic crust. Using the stratigraphy ofthe classic Penrose ophiolite and more modernestimates of unit thicknesses, our standard section(Figure 2) has a total crustal thickness of 7.1 km(White et al., 1992; C. Z. Mutter and J. C. Mutter,1993), which includes 1,000 m of extrusives,1,100 m of sheeted dikes, and 5,100 m of gabbros.This “standard” section translates into a volumeproduction rate of 24.14 km3 yr21 using Parsons’(1984) estimate of the ocean-crust surface area

production rate of 3.4 km2 yr21. This rate ofoceanic-crust production can be translated into amass production rate by using in situ densitiesmeasured in the oceanic crust. These values rangefrom ,2,700 kg m23 for extrusives (Salisburyet al., 1979), to ,3,000 kg m23 for gabbros(Dick et al., 2000). The sheeted dikes areextrapolated to have an average density of2,850 kg m23. This yields a total ocean-crustproduction rate of 7.05 £ 1016 g yr21. However,it must be emphasized that there are still significantuncertainties in these total fluxes and productionrates.

The Troodos and Samail ophiolites (Figure 1)contain all the essential components of ourstandard section in Figure 2. Both of these

300 0 60 120 180

332B 417A,D 418A

504B

543A

735B

765C

801C 896C

Troodos ophioliteSamail ophiolite

Figure 1 Locations of the drill sites and ophiolites discussed in the text.

Table 1 Ophiolite and ocean-crust drill hole locations.

Site Lat. Lon. Age(Ma)

Ocean/Region Depth(m)

Recovery(%)

Comments

332B 36.945 33.641 3.5 Atlantic 580 21 W MAR417A 25.106 268.041 120 Atlantic 206 70 Highly altered 150 m

abyssal hill417D 25.106 268.041 120 Atlantic 366 70 Normal crust418A 25.035 268.057 120 Atlantic 544 72 Normal crust504B 1.227 283.734 5.9 E. Pacific 2,000 20 MORB extrusives/dikes543A 15.712 258.654 82 Atlantic 44 81 MORB extrusives735B 232.722 57.264 9 Indian Ocean 1,000 MORB gabbros/troctolites765C 215.976 117.592 155 Indian Ocean 247 31 Normal MORB extrusives801C 18.642 156.360 157 W. Pacific 135 60 Alkalic basalts and normal

MORB extrusives896C 1.217 283.723 5.9 E. Pacific 290 28 Normal MORB extrusivesTroodos 34.800 32.000CY 1 35.048 33.179 92 Tethys/Cyprus 1,150 95 Suprasubduction zone

extrusivesCY 1ACY 4 92 Tethys 2,300 95 Dikes gabbros, ultramaficsSamail 24.500 58.000 120 Tehtys/Oman Normal MORB

All locations given in decimal degrees, negative values give southern latitudes and western longitudes.

The Unaltered Oceanic-crust Protolith 513

ophiolites were formed in the ancient TethysOcean (92 Ma for the Troodos Ophiolite and 100–130 Ma for the Samail ophiolite). The Samailophiolite represents crust that is quite similar to“normal” oceanic crust (Kelemen et al., 1997), butit has not been studied for geochemical massbalances. The Troodos was studied extensively forits primary and hydrothermal characteristics but itreflects an oceanic crust that is more typical ofsupra-subduction zone settings, having geochemi-cal characteristics distinct from normal mid-oceanridge basalt (MORB) (Robinson et al., 1983;Schmincke et al., 1983). For this reason, the bulkcomposition of the Troodos ophiolite is not typicalof average oceanic crust, even though it is clearlya part of the spectrum that is produced andrecycled. The Troodos ophiolite is well studied,partly because of an intense drilling campaign thatyielded a total of 8,000 m of core material from allstructural levels in the crust. Key drill sites includeCY1 and CY1A, with a total penetration of1,175 m though extrusives (Gibson et al.,1991a), and CY 4, that penetrated 2.3 kmof sheeted dikes, gabbros, and pyroxenites

(Gibson et al., 1991b). Their stratigraphic positionis plotted in the reference section in Figure 2.Recovery from all of these drill cores was found tobe ,100%.

There are a large number of drill sites in oceaniccrust, but four sites stand out for their impact onstudies of hydrothermal alteration processes.These include sites 417 and 418, which werestudied to constrain the chemical mass fluxes ofthe upper crust, site 504B for the mid crust orsheeted dikes and site 735B for the gabbro section(Figures 1 and 2; Table 1).

Three drill holes at sites 417 and 418 probeupper oceanic crust of the western Atlantic. Thiscrust formed at the relatively slowly spreadingmid-Atlantic ridge, which may be more fracturedthan fast spreading crust, and contain morepillows and less massive flows. Site 417A, wasdrilled into a 150 m high abyssal hill with a totalpenetration of 208 m, whereas sites 417D and418A were drilled in the nearby abyssal plain, withpenetrations of 360 m and 544 m, respectively.Recovery rate was about 70% at all of thesites. Site 417A contains, on average, 20%

Top

1,000

2,000

3,000

4,000

5,000

6,000

7,000

Sheeted dikes1,100 m / 2,850kg m–3 1.06 × 1016 g yr–1

5,000m /3,000 kg m–3

5.1×1016g yr–1

Upper extrusives(600 m)

Lower extrusives(400 m)

Sheeted dikes(1,000 m)

Gabbros(3,250 m)

Gabbro-norites(1,000 m)

Ultramafics(750 m)

Gossansmassive sulfides

Spilites

Epidosites

Plagiogranites,amphibolites

Extreme alteration:

Serpentinites,chromitites, talcum

418A 504B

735B

Total7,100 m 7.08 × 1016 g yr–1

CY 1CY1A

CY4

Gabbrosultramafics

Extrusives1,000 m /2,700 kg m–3

9.18 × 1015 g yr–1

Figure 2 The ocean crust reference section used in this paper, using a standard “Penrose” style ophioliteassemblage (Penrose Conference Participants, 1972) and a crustal thickness after White et al. (1992) and C. Z. Mutterand J. C. Mutter (1993). Various modeling parameters used are also indicated, including densities and fluxes, theapproximate positions of drill holes discussed in this paper, and some typical extreme alteration environments found

at various depths of the oceanic crust.


volcaniclastics, whereas sites 417D and 418Acontain only ,6% volcaniclastics (Robinson et al.,1979). The higher abundance of volcaniclasticsand the more exposed nature of Site 417A resultedin more intense alteration than at sites 417D and418A (Donnelly et al., 1979a,b,c). Lithologies atSites 417D and 418A are quite similar, but thecrust was formed at different mid-ocean ridgevolcanoes, offering a comparison between twoseparate but quite similar crustal sections.Staudigel et al. (1995, 1996) combined thesesites into one representative section of almost600 m length, including 20% materials from 417Aand 40% each from 417D and 418A in the upper150 m, all adjusted to 6% volcaniclastics.

Additional sites that sample upper oceanic crustinclude 332B, 504B, and 801C (Figure 1; Table 1).Site 332B was also drilled in the Atlantic on veryyoung crust, and 504B and 801C were drilled inthe eastern and western Pacific, respectively.Amongst those, site 504B is the deepest drillhole into ocean crust, sampling both extrusivesand the underlying sheeted dikes, thus offering aunique opportunity to study alteration processesin the deeper reaches of the crust (Alt et al.,1993a,b). Site 801C is located in the westernPacific, on 170 Myr old oceanic crust andpenetrated almost 500 m of basalt with averagerecovery rate of 47% (Plank et al., 2000). Thishole recovered alkalic basalts near the top,overlying a substantial section of “normal”oceanic crust below. The deeper portion of thishole is still under investigation.

Samples of the deepest oceanic crust areaccessible in only one ocean drill core: site735B was drilled on 11 Myr old crust on the SWIndian Ridge (Dick et al., 2000). This holepenetrated a few tens of meters of pillow lavasand ,1,200 m of gabbros, with ,100% recovery.The site never reached cumulates (Dick et al.,2000). The lack of sheeted dikes and the nearabsence of pillow lavas is clearly quite differentfrom the “normal” ophiolitic crust, but site 735rock types are not very different from materialsfound in the upper plutonic section of ophiolites.Thus, site 735B is used to represent the deepercrust for the composite crust section.

3.15.2.2 Estimating Unaltered Oceanic CrustCompositions

Estimating the unaltered composition of theoceanic crust is necessary for flux estimatesbecause these estimates are determined as thedifference between original and altered compo-sitions. Table 2 displays an average unalteredmid-ocean ridge basalt composition (MORB,Hofmann, 1988) and estimate of an unfractionatedMORB composition (Sun and McDonough, 1989),but it is important to emphasize that the

composition of MORB can vary substantially(e.g., Chapter 3.13 or http://petdb.ldeo.columbia.edu; Lehnert et al., 2000). This variation is toolarge for most elements to allow the use of a singlefresh “average MORB” as a starting compositionin chemical flux estimates. Thus, fresh compo-sitions have to be constrained for any particularsuite of altered basalts studied.

There are two main ways in which the “fresh-rock” composition of an altered rock is deter-mined. One is a petrographic – geochemicalapproach, which involves re-assembling a rockfrom the chemical analyses of residual igneousphases still present in the rock using their originalmodal abundances. Another is a purely geochemi-cal approach that uses the chemical systematics infresh rocks and placement of an altered rock insuch fresh-rock systematics using alterationinsensitive parameters. Deviations from theextrapolated fresh-rock composition are theninterpreted as being caused by alteration pro-cesses. In both methods, there are significanterrors. Interpretations of chemical trends have tocontend with the natural scatter of sampledistributions in correlation diagrams, as well asthe cumulative analytical errors introduced fromregressions, normalizations, and analytical uncer-tainty. Petrographic–geochemical reconstructionsare affected by small-scale heterogeneities inmineral distribution and uncertainties in modalestimates. These errors are especially large formajor elements or trace elements that arecompatible in olivine and plagioclase, includingmagnesium, calcium, silicon, sodium, aluminum,iron, nickel, and barium. These errors are thelowest for incompatible elements that are presentin very low abundances in the original igneousrock, such as H2O, CO2, uranium, potassium,rubidium, caesium, REEs, and thorium.

A third approach that is commonly used toconstrain chemical fluxes compares differentlyaltered materials, such as altered pillow marginsand less altered pillow interiors, or samples withor without alteration haloes around veins (e.g., Altet al., 1986), mineralized and unmineralized zonesor differently altered gabbros (e.g., Bach et al.,2001) in order to constrain chemical changesassociated with alteration. However, “least”altered samples only rarely reflect the originalcomposition reliably. A second problem in thisapproach is the relatively small sample sizestypically analyzed from ocean drilling materials.Typical sample sizes are about 15 cm3, which issmall when compared with local variability inmodal mineralogy. Indeed, individual phenocrystphases can be several millimeters in size. Localvariability in modal mineralogy is particularlycommon in pillow lavas where phenocryst abun-dances can vary as a function of radial distancefrom the center or vertically within the center

The Unaltered Oceanic-crust Protolith 515

http://petdb.ldeo.columbia.edu


Table 2 Hydrothermal fluxes.

Fresh

MORB

Units Bulk rock gains (þ) and losses

(2) for extrusives

Dikes Gabbros Average crustal

gains/losses

Units Total flux

from crust

Units Hydrothermal fluxes from submarine vents

(g yr21)

River fluxes

(g yr21)

0–600 m Error 600–1000 Flank Axis low Axis high

SiO2 50.45 wt.% 1.18 1.5 0.5 0.25 0.1 0.23704225 g/100g 21.67E þ 14 g yr21 1.08E þ 12 1.80E þ 13 6.61E þ 13 3.85E þ 14

Al2O3 15.26 wt.% 1.94 3 29.65E þ 13 g yr21

FeOtot 10.43 wt.% 20.01 0.03 0.00E þ 00 1.65E þ 12 1.37E þ 13 1.65E þ 12

MnO 0.19 wt.% 0 0.01 0.75 20.035 3.69E þ 10 4.97E þ 11 3.97E þ 12 3.48E þ 11

MgO 7.58 wt.% 20.19 1 0 0 20.2 20.15690141 g/100 g 1.11E þ 14 g yr21 22.18E þ 14 24.03E þ 13 21.05E þ 14 2.18E þ 14

CaO 11.3 wt.% 2.14 1.5 0.3 0.04 20.1 0.13352113 g/100 g 29.43E þ 13 g yr21 2.64E þ 14 3.03E þ 11 1.23E þ 14 6.73E þ 14

Na2O 2.679 wt.% 0.15 0.2 0.15 0.15 0.03 0.06549296 g/100 g 24.63E þ 13 g yr21 23.84E þ 13 3.53E þ 14

K2O 0.11 wt.% 0.54 0.15 0.05 0 0 0.0484507 g/100 g 21.71E þ 13 g yr21 23.11E þ 13 6.12E þ 12 3.01E þ 13 1.22E þ 14

Rb 1.26 ppm 10.3 3 1 0 0 0.92676056 mg/kg 26.55E þ 09 g yr21 22.22E þ 09 1.37E þ 10 2.39E þ 10 3.16E þ 10

Cs 0.0141 ppm 0.183 0.050 0.03 0 0 0.01715493 mg/kg 21.21E þ 08 g yr21 6.40E þ 08

CO2 0.15 wt.% 3.26 0.5 0.5 0.06 0.06 0.35521127 g/100 g 22.51E þ 14 g yr21 27.04E þ 12 21.76E þ 12 27.04E þ 12 1.41E þ 15

H2O 0.2 wt.% 2.81 1 2.09 1.5 0.11 0.44873239 g/100 g 23.17E þ 14 g yr21

S 960 ppm 20.06 20.06 20.06 0.02 20.0037 g/100 g 2.59E þ 12 g yr21 23.53E þ 13 21.60E þ 13 24.49E þ 13 2.85E þ 13

Li 4.5 ppm 2.8 2.8 20.5 22 22 mg/kg 7.71E þ 09 g yr21 21.25E þ 10 4.8587E þ 10 4.37E þ 11 9.72E þ 10

B 0.5 ppm 25.7 10 4 5.6 2.2 4.81 mg/kg 23.40E þ 11 g yr21 1.84E þ 11 6.49E þ 09 7.89E þ 10 5.84E þ 11

Sr 113 ppm 22 5 3 0.3 21 1.4 mg/kg 29.68E þ 10 g yr21 2.19E þ 11 29.00E þ 08 4.60E þ 09 2.02E þ 11

U 0.0711 ppm 0.3 0.05 0.1 0.01 0.007 0.03746479 mg/kg 22.65E þ 09 g yr21 9.60E þ 09

0.3 23.60E þ 14 g yr21 1.90E þ 14 21.14E þ 13 1.02E þ 14 2.23E þ 15

of a pillow. Staudigel et al. (1996) estimatederrors in chemical composition on the basis ofvariation in phenocryst modal abundances andfound relatively large errors for some elements,particularly MgO (e.g., 7.7 ^ 0.2 wt.%).

One of the key objectives in estimating freshrock compositions is to determine the chemicalinventory present before alteration in a givenreference volume of altered rock. This is differentfrom the “fresh-rock composition,” becausewater–rock interaction occurs in an open system.The pitfalls arising from open-system behaviorcan be illustrated in two examples: secondarymineral precipitation and rock dissolution.

In rocks with significant pore space, mineralsmay precipitate in these pore spaces duringalteration. For example, precipitation of 5 g per100 g of calcium carbonate in the pore spaces of afresh rock dilutes the abundances of all otherchemical components by 5%. In this example,carbonate precipitation causes an apparent loss of,2.5 wt.% SiO2, when in fact SiO2 remainedimmobile during the alteration.

The opposite effect may be caused by wholesale(congruent) dissolution of the rock wherebymassive amounts of basalt could be lost from agiven rock volume, without any trace of concen-tration change. However, truly congruent dissolu-tion is quite unlikely; dissolution typically leavesbehind some residual immobile material.Titanium is generally assumed to be immobile,and the abundance of titanium in a rock cantherefore be used to quantify open-systemchemical behavior. The first step in addressingopen-system behavior is to estimate the titaniumconcentration in the fresh rock equivalent, whichis taken as the mass of titanium originally presentin a rock per 100 g of the material analyzed.The fraction of titanium originally present,together with an independent estimate of thefresh rock composition, allows an estimation ofthe concentration of all the other elements in thefresh rock. Chemical fluxes, then, are calculated asthe differences between elemental concentrationin the unaltered rock (per 100 g analyzed) and theconcentration of the element in the alteredmaterial. The most important sources ofuncertainties in this procedure arise from theuncertainties in titanium determination, errors inthe estimate of the titanium concentration in theunaltered rock (e.g., Staudigel et al., 1995, 1996)and errors due to small-scale titanium variationswhen using the titanium from a “fresh”-alteredsample pair (e.g., Bach et al., 2001).

It is possible to estimate the validity of theconstant titanium assumption through a compari-son with other, independent methods, such asworking with constant reference volumes andestimating the amount of initial rock presentbased on density and pore space consideration.

Specifically, if the density is known for the freshand the altered rocks, fresh and altered compo-sition can be expressed on a volume percent basis,and fluxes are calculated relative to a constantvolume. This method was evaluated at sites 417and 418 and shown to produce almost identicalresults as the constant titanium assumption(Staudigel et al., 1996).

3.15.3 DETERMINING THE ALTEREDCOMPOSITION OF THE OCEANICCRUST

Determining the composition of altered oceaniccrust is also not very straightforward. Several keysteps are required in obtaining a meaningfulestimate of altered rock compositions. First, thechoice of a study site is important. Is the sectionold enough to have experienced the bulk of itsalteration processes? Are the alteration patternsrepresentative of most oceanic crust, or are theyonly of local importance? Is core recoverysufficient to allow a representative estimate ofthe section drilled? Once these three questions areanswered positively, a method has to be workedout to determine the bulk composition of the crustand chemical fluxes on scale lengths that aremeaningful for global chemical budgets. Thefollowing sections evaluate problems related tothese questions, in particular the role of recovery,the range of “typical” types of alteration, con-straints on the duration of alteration in the crustand techniques for determining representativecompositions of a heterogeneous medium. Theseproblems are illustrated with examples fromparticular ocean drill sites or ophiolites.

3.15.3.1 Recovery Rate

It is extremely rare in ocean drilling to have100% recovery of basement materials. Most drillcores from basement consist of variously roundedfragments that typically don’t match up with theneighboring core fragments. Recovery rates arelow, particularly in the upper and young oceaniccrust, which has not been effectively sealedthrough mineral precipitation and fractured onthe same scale length as the drill bit (and smaller).Competent rock fragments, resistant to the cuttingaction of a drill bit, are held together by relativelysoft materials like clays, carbonates, or chlorite.The stress imposed on resistant rock fragment islikely to break up the loosely cemented formation,and grind up the soft material filling the void andfracture spaces in between. This material is thenejected from the hole with the drilling fluid on theocean floor as mud, sand, or chips and not in thecore barrel. This problem also affects the innerwalls of the drill hole, where competent rocks are

Determining the Altered Composition of the Oceanic Crust 517

more likely to be exposed than soft rocks. Mostalteration materials are not found in distincthorizontal layers, but occur as pockets betweenor within pillows, or as veins that commonly cutthe hole at very steep angles, which can becompletely eroded, even though the neighboringmaterial is actually recovered. A similar samplingbias occurs in outcrops of ophiolite sections,where clays are much more likely to be erodedthan fresh, well-cemented basalt. These problemscan be overcome only by complete recovery ofdrill core.

The recovery rates of various ocean-crust drillholes are given in Table 1. Recovery rates aredetermined as the ratio of the cumulative length ofrecovered fragments to the actual length of thesection drilled. This estimate is quite meaningfulat high recovery rates, but not at low recoveryrates, where individual pieces tend to be roundedoff and have a smaller diameter, resulting involume recovery rates that are substantiallysmaller than the linear rates. Thus, low recoveryrates practically eliminate substantial portions of adrill hole from direct study and they can make itvery difficult, if not impossible, to estimate bulkcompositions from bulk rock data, particularly inpillow sections and breccia zones. Drilling inoceanic crust often has very low recovery, on theaverage of 25–30%, with some sections havingnearly zero recovery (e.g., a substantial fraction of504B). At an average recovery rate of 30%,proportionately more material is recovered frommassive units (50–100%), than breccia zones(0–10%) complicating the estimates of alterationinventories. Ocean drill sites with the highestrecovery rates include site 801C, with almost 50%recovery, 417A, 417D, and 418A, with over70% recovery, and 735, with 86% recovery.The Cyprus drill cores CY 1 and CY 4 havenearly 100% recovery. The other deep drillsites discussed here, 504B and 332B havesubstantially lower recovery rates, on the orderof ,20–30%.

3.15.3.2 Types of Alteration

One of the key goals in quantifying alteration isto determine the respective contributions of

different types of alteration. The focus of theseefforts have to be on what is considered “normal”or “most representative,” but it is also important toexplore the complete range of alteration behaviorin the oceanic crust. Even unusual alterationenvironments may have a significant impact onthe geochemical behavior of some elements, andthey may cause distinct geochemical behavior insubduction zones or in the mantle. For this reason,it is important to focus on the “average”geochemical behavior, but keep an eye on thecompositional diversity as well.

Studies of ophiolites have been particularlyuseful in identifying types of alteration in oceaniccrust and for understanding their relative signifi-cance. Many of these alteration features have alsobeen found in ocean drill cores, but many aspectsof seafloor hydrothermal alteration remainunexplored by ocean drilling.

The upper extrusive oceanic crust (0–600 m,Figure 2) is primarily altered at low temperatures(,100 8C). Alteration is commonly not pervasive,whereby igneous phases (glass, phenocrysts) maycoexist with alteration phases (clays, zeolites, andcarbonates). At high water–rock ratios, oxidativemineral assemblages (e.g., celadonitic clays) formintensely altered zones, but most ocean-crustalteration is accomplished by more reducing fluidsat lower fluid–rock ratios (Table 3). However,in most low-temperature altered basalts, highlyreducing minerals (like pyrite) may coexistin close proximity with oxidizing minerals likeceladonite or hematite. The distinctive differencesbetween mostly oxidized and reduced alterationenvironments are well illustrated by compar-ing the oxidative upper (abyssal-hill) portionof site 417A to the more reduced sectionsat sites 417D and 418A (e.g., Donnelly et al.,1979a,b,c; Alt and Honnorez, 1984). A similarcontrast exists between Cyprus drill cores CY 1and CY 1A, whereby CY 1, from the uppermostAkaki canyon, closely resembles the oxidizedalteration in site 417A and the remaining coreshows alteration behavior more like site 417D and418A (Gillis and Robinson, 1988; Bednarz andSchmincke, 1989). Based on the distribution ofalteration features in the northern portion of theTroodos ophiolite, Staudigel et al. (1995)

Table 3 Secondary minerals in the oceanic crust.

Upper extrusive crust Deeper extrusives

Oxidizing conditions Nonoxidizing

Aragonite, analcite, calcite,celadonite, chalcedonyFe-hydroxide, hematite,philipiste, K-feldspar,saponite

Anhydrite, analcite, calcite,celadonite, Fe-hydroxide,mixed layer chlorite–smectite,Na zeolite pyrite, saponite

Albite, calcite, chlorite,epidote, pumpellyite,prehnite, quartz, sphene


suggested that the 417A style of oxidativealteration makes up ,20% ( 10%) of the upper150 m of oceanic crust, while the reduced styleof alteration seen at sites 417D and 418A arerepresentative of 80%.

The upper oceanic crust can show extremecompositional variation near hydrothermal vents.Based on the abundance of hydrothermal depositsin the Troodos ophiolite, and the frequency ofblack-smoker-type deposits in mid-ocean ridges,these highly mineralized zones are likely torepresent less than one-tenth of a percent of thevolume of the total upper oceanic crust. Thus formost elements, these deposits are not veryimportant in the total mass balance. The exceptionto this is elements that are enriched in ventdeposits by several orders of magnitude relative tobasalt (e.g., copper, zinc, lead, manganese, iron,nickel, cobalt, platinum, silver, gold; see alsoChapter 3.12). This limits the significance of thesedeposits to trace-element mass balances in theoceanic crust. A convincing case for this wasmade by Peuker-Ehrenbrink et al. (1994), whosuggested that global lead cycles in seawatermight be substantially influenced by the precipi-tation of a small amount of hydrothermalmetalliferous sediments on top of the oceaniccrust.

In the deeper extrusive oceanic crust and in thesheeted dikes, alteration temperatures increase(.100 8C), water–rock ratios decrease, and mostprimary igneous phases tend to be almost entirelyreplaced by secondary phases (the exception againis near hydrothermal conduits, where temperatureand water–rock ratios are high). In this depthrange glass, olivine, and calcic plagioclase aretypically replaced by greenschist-facies mineralassemblages (Table 3). Typical phase assem-blages in deep extrusives and dikes for in situoceanic crust are best described at site 504B(Alt et al., 1986, 1993, 1996), offering importantinsights into alteration processes in this depthrange.

Several important alteration environments,however, are not observed for in situ oceaniccrust. These include spilites (Cann, 1969), a rocktype that displays almost complete exchange ofcalcium for sodium, leading to formation of analbite-rich rock, and epidosites, the metal depletedepidote–quartz–chlorite assemblage that is likelyto be characteristic of the reaction zones, or atleast last equilibration zones of black-smokerfluids (e.g., Schiffman and Smith, 1988;Richardson et al., 1987; Bettison-Varga et al.,1992). In particular, the importance of spilites inchemical mass balances could be large, but theyare effectively unknown for in situ oceanic crust.Spilites are therefore not considered in any massbalances. Epidosites are likely to be important forthe mass balance of some trace metals in the

oceans, but are unlikely to influence the oceancrust generation – subduction budget becausethey are probably just as uncommon as themassive sulfide deposits. Another potentiallyimportant alteration assemblage is predictedfrom experimental studies and observations fromblack-smoker chemical compositions. Heatingseawater to 450 8C results in massive precipitationof sulfate (mostly anhydrite; Bischoff andSeyfried, 1978) and may result in boiling, leavingbehind brines (Butterfield et al., 1990). Anhydriteand brines are likely to be stored only temporarilyin the oceanic crust and will be dissolved byseawater circulating through the crust at a latertime after much of the magmatic heat isexchanged with seawater (e.g., Alt, 1994, 1995).

Deeper levels in the oceanic crust displayhigher-grade hydrothermal alteration, rangingfrom amphibolite grade to anatexis (Table 3),where the formation of plagiogranites has beenassociated with the partial ingestion ofhydrothermally altered ocean crust materials.Hydrothermal alteration processes in this regimehave been studied in materials recovered fromoceanic fracture zones, particularly in site 735B(Dick et al., 2000), as well as in ophiolites (Nehliget al., 1994). Overall, it appears that the upper200 m section of site 735B displays unusuallyhigh water–rock ratios due to its proximity toseawater, but the bulk of this 1.5 km section aremore typical of normal crust (Dick et al., 2000).However, 735B does not display any of theextreme varieties of alteration observed in someophiolites or in near-surface exposures on theocean floor, such as serpentinites, or other hydrousequilibrium assemblages such as talcum deposits(e.g., Bonatti, 1976; Baer, 1963). Such compo-sitional domains are likely to be formed atrelatively slow spreading centers where brittledeformation may carry water to great depth,deeper than in the faster-spreading oceanic crust.Such deposits could potentially be very importantfor mass balances, particularly for recycling ofwater, but are not included in this review due tolack of data.

3.15.3.3 Duration of Alteration

Determining that alteration is complete or nearcomplete is one of the major pre-requisites fordetermining alteration fluxes at any study site, andfor this reason, it is important to determinethe duration of alteration in the oceanic crust.There are two main constraints for this: measure-ments of heat flow, which indicate the durationof convective heat loss, and isotopic dating ofsecondary minerals that precipitate duringalteration of the oceanic crust.

Heat-flow measurements can be used to deter-mine the duration of seafloor alteration by


comparing the absolute estimate of heat flow toheat flow calculated from plate-cooling models.The difference between the two is then assigned toconvective heat flow. The difference betweentheoretical and measured heat flow suggests totalconvective heat loss of approximately 11 TW(Stein et al., 1995). About one-third of this heatloss occurs within the first million years after crustformation, the second one-third occurs between1 Ma and 8 Ma, and the last one-third occursbetween 8 Ma and 65 Ma (Stein et al., 1995;Figure 3). However, the “termination” of con-vective heat loss at 65 Ma also coincides with thetime beyond which the simple relationshipbetween ocean-floor depth and plate-coolingmodels break down. Thus, a comparison ofmeasured and theoretical heat flow is not verymeaningful at this age.

Alternatively, the duration of hydrothermalconvection in the oceanic crust can be estimatedby mapping the distribution of nonlinear tempera-ture profiles taken during heat-flow measurementsas a function of oceanic-crustal age. Purelyconductive heat loss (i.e., no hydrothermalcirculation) results in linear temperature profilesin sediments, while convective heat loss results inconcave or convex profiles, depending on whetherthe water penetrates into or comes out of thesediments.

Both heat-flow methods deliver similar results,suggesting that convective heat loss, hence hydro-thermal alteration, may occur for time periods

of up to ,65 Ma (Parsons and Sclater, 1977;Langseth et al., 1988; Anderson et al., 1979; Steinet al., 1995; Pelayo et al., 1994).

The duration of chemical exchange betweenseawater and basalt can also be determined byisotopic studies of hydrothermal minerals.A variety of techniques have been used, rangingfrom direct dating by K/Ar and Rb/Sr isochrontechniques to comparisons of the initial strontiumisotopic composition of alteration minerals withthe isotopic evolution of seawater (Gallahan andDuncan, 1994; Richardson et al., 1980).

Compiled isochron data for the duration ofalteration at a series of DSDP and ODP sites areshown in Figure 3. The ages thus calculated arethe differences between crustal age and vein-mineral age. Due to the rather large spread ofdates, ages were binned into 20 Ma intervals, withaverages at the mid point of each interval shown inFigure 3. Even though rather crude, this distri-bution shows a remarkable resemblance to theheat-loss curve. While the resolution of samplingdoes not allow any strong constraints on the actualage distribution in the (most critical) first 20 Ma,they do show very clearly that vein-mineralprecipitation can continue for a rather long timein this global distribution of sampling sites.Collectively, these observations suggest thatoceanic crust younger than 10 Ma is unlikely tohave experienced the complete cycle of seaflooralteration.

3.15.3.4 Determining the Composition ofExtremely HeterogeneousAltered Crust

Oceanic-crust alteration involves formation ofdistinct compositional domains, which range insize from a ridge segment to a submilimeter sizedvesicle filling. Obtaining a robust average of suchchemically heterogeneous material is critical forunderstanding the chemical fluxes associated withseafloor alteration. Averaging heterogeneitieswith length scales smaller than the size of atypical geochemical sample is relatively easy, butbecomes more difficult as the size of theheterogeneity increases. For large-scale hetero-geneities, multiple samples need to be taken andcharacterized with respect to the proportion eachsample contributes to the average. Such samplesmay be mixed in the correct proportions toproduce composite samples that represent anaverage for a given section or crust type, or theymay all be analyzed individually and thenaveraged arithmetically.

Large-scale geochemical sampling requires astrategy that bridges the gap between sample sizeand the scales of key heterogeneities. Hetero-geneities on the order of several meters to tens

20

40

60

80

100

Fraction of total convective heat loss

% vein mineral ages

0 20 40 60 80Million years

Figure 3 Cumulative convective heat loss (Booijet al., 1995) of the oceanic crust and completion of veinmineral deposition (Stein et al., 1995) in the oceaniccrust. Heat flow curve after C. A. Stein and S. Stein(1994b), including vein mineral ages from site 261 afterHart and Staudigel (1986) site 417/418 after Hart andStaudigel (1978, 1986) and Richardson et al. (1980),site 462A after Hart and Staudigel (1986) site 516 (the18 Ma Rio Grande Rise), and 597 (South Pacific) afterHart and Staudigel (1978, 1986) and the TroodosOphiolite after Staudigel et al. (1986), Booij et al.(1995) and Gallahan and Duncan (1994). Vein mineralage cumulative curve is binned in 20 Myr age groups.


of meters are probably the smallest relevantlarge-scale heterogeneities that need to be avera-ged. A compositional domain of this size isrelevant geochemically because it may reflect asingle chemical system, with its own charac-teristic phase assemblages that display consistentbehavior during dehydration or melting events.This length scale is also a practical size foranalyses of compositional domains through localobservation of continuous exposure or coresections. Data from such a length scale can berelatively easily extrapolated to larger scales byvisual integration of (.kilometer scale) cores, orfield outcrop areas.

The selection of individual samples for large-scale compositional study may follow differentstrategies, depending on the type of crust studiedand personal preferences. One of the majordifferences in strategy is the use of compositesamples, versus the analysis of individual samples.Procedurally, these two approaches are identical,as long as the sampling strategy is comprehensiveand the modal abundance data are determined in aconsistent fashion. The two approaches each haveadvantages and disadvantages.

Analysis of individual samples has the advan-tage that it offers the potential for resolving thechemical properties of different alteration types,before the samples are combined arithmeticallyinto one average. The disadvantage of thisapproach is that it requires a large number ofanalyses and may be too labor intensive for somegeochemical parameters. The advantage ofcomposite samples is the significantly lowernumber of analyses required, which makes itmuch more likely to obtain a completegeochemical characterization of a representativeset of samples at a particular crustal section.

However, it is also possible to find a compromisebetween the use of composite or the individualsamples. Staudigel et al. (1995, 1996) mixed one“Super” composite that represents the average ofthree drill sites, but they also analyzed two types of“subcomposites” to understand uncertainties andthe internal structure of the “Super” composite.One type of subcomposite contrasts thecompositions of volcaniclastics with flows thatallow to contrast extreme with moderately alteredregions. In addition, they made independent“depth” composites for two sites with similaralteration types to obtain an estimate of howvariable such estimates can be for differentsections. These subcomposites provide someinsights into the chemical makeup of a site, withouthaving to analyze all samples individually.Alternatively, it is possible to analyze all individualsamples for geochemical parameters that arerelatively easy to determine, and using compositesonly for the more involved analytical steps.The latter was done for all the highly altered

volcaniclastic samples in the composites from sites417 and 418 (Staudigel et al., 1996).

Strategies for selection of individual samplesmay also vary. One approach is to use relativelylarge, representative samples that include veins,vugs, and variously altered haloes or othermaterials (Staudigel et al., 1989, 1995, 1996).Because of the complexity of alteration and thelarge number of permutations of primary andsecondary features, several samples are typicallyrequired to approximate the chemical inventory ofa particular compositional domain. This approachhas the benefit of averaging out local heterogenei-ties or chemical gradients between veins and hostrock, and minimizing uncertainties and contami-nation problems during sample handling. Anotherapproach is to separate all major alteration typesand vein materials, study them individually andrecombine them in their respective proportions(e.g., Alt and Teagle, 1999; Bach et al., 2001).The advantage of this method is that it sheds lighton the alteration processes and their impact onchemical fluxes more fully than composite samplesdo. The disadvantage is that it produces largeruncertainties in the overall averages, and requiressubstantially more sample handling and analyses.

Any attempt to produce a bulk compositionalestimate of altered ocean crust is criticallydependent on accurate determination of theproportions of different rock types in the volumeof interest. Simply averaging all analyses pub-lished for a particular drill site does not yield arealistic bulk compositional estimate for tworeasons. First, low recovery during drilling biasesthe average towards the least altered material,whereas high recovery rates include a greaterproportion of altered material. A hypothetical caseof several holes drilled into the same crust atdifferent recovery rates will show a positivecorrelation between degree of alteration andrecovery rate. Secondly, the choice of samplesfor a particular study depends strongly on thescientific goals of that study. Sampling for igneousgeochemistry studies typically focuses on the leastaltered samples. Such studies often further bias thesample selection by crushing the rocks andselecting the freshest rock chips out of a samplefor analysis, and leaching carbonates out ofthe rock to obtain a maximally pristine chemicalcomposition of an igneous rock. Such studies areobviously not useful in constraining the amount ofalteration at a site. On the other hand, studies ofalteration mineralogy of the crust focus theirsampling on highly altered sections, especiallyextreme alteration types. Thus, a simple averageof published geochemical data from a particularsite is likely to offer insights mostly into therecovery rate and the types of studies performedrather than the actual average composition ofoceanic crust at a particular site.


3.15.4 CHEMICAL CHANGES IN ALTEREDCRUST COMPOSITION DUE TOHYDROTHERMAL PROCESSES

3.15.4.1 Time Dependence of Crust Hydrationand Carbonate Addition

Uptake of water and CO2 is one of the mostsensitive indicators of alteration, and producessome of the most profound chemical changes inaltered oceanic crust. These chemical changes arecumulative and offer the opportunity to evaluatethe age dependency of alteration in drill sites ofvarious ages. Presently, there are three sites in theupper oceanic crust that can be used for suchcomparisons: the 3.5 Myr old DSDP site 332B,the 5.5 Myr old site 504B, and the 120 Myr oldsites 417A, 417D, and 418A (Figures 4(a) and (b)).Sites 417 and 418 are in oceanic crust that issubstantially older than the predicted duration ofhydrothermal alteration derived from heat-flowdata (Figure 3) and thus these sites representmature oceanic crust, which has experienced thecomplete history of hydrothermal alteration. Crustpenetrated in site 504B is much younger than thatin sites 417/418, and is in the age range of crustthat is expected to have active hydrothermalalteration, based on the global heat-flow data set,even though local heat-flow measurementssuggest that alteration is completed at this site

(Langseth et al., 1988). Crust of site 332B is evencloser to the mid-ocean ridge and is expected to beexperiencing active hydrothermal circulation.Unfortunately, sites 332B and 504B have verylow recovery rates, and for this reason, some effortis needed to determine the true differences intime-variant alteration behavior.

The highest contents of H2O are found in theupper 300 m of crust from sites 417A, 417D, and418A (Figure 4(a)). At the same depth, theyoungest crust at site 332B shows the leastintense hydration, while crust at 504B hasintermediate water content. Water contents ofcrust from 417/418 decrease with depth, whilewater contents at 332B and 504B increase slightlywith depth before decreasing again. All trendsconverge at depths greater than 300 m. Differencesin recovery rates is an unlikely cause for the trendsobserved in the upper 500 m of this diagram, and itis most likely that 332B and 504B have not reacheda mature degree of hydration in their upper portion,while 504B is likely to have completed itshydration at depths .600 m. Thus, the formationof layer silicates in the upper portion of 504B and332B is probably an ongoing process.

CO2 shows behavior similar to water. Itsconcentration is also highest in the upper 300 mof crust at sites 417 and 418 and lowest in 504Bthroughout the depth intervals covered by allholes. It is interesting to note that crust of site

00 22 66 8

800

600

200

400

1,000

44wt.% CO2

504B332B417 / 418

Super417D418A

wt.% H2O

0

(a) (b)

Dep

th (

m)

Figure 4 (a) Water versus depth and (b) CO2 versus depth for major deep drill sites in the Atlantic and PacificOceans. Data compilations are taken from the Initial Report Volumes of the Deep Sea Drilling Project orthe Ocean Drilling Project (site 332B: Aumento and Melson, 1977; 417/418 Donnelly et al., 1979b; site 504B:Alt et al., 1993; bold lines connect depth composite data from sites 417D and 418A (after Staudigel et al., 1995).


332B has higher CO2 values than that of 504B,particularly in the depth interval 400–500 m, eventhough it samples younger crust. With increasingdepth, CO2 values of 417/418 crust show asystematic decrease to values that are substantiallyabove the values from site 504B for some depth.These observations indicate that CO2 uptake inoceanic crust is highly variable in younger crust,and both young crustal sections are distinct frommature ocean. Site 504B appears to have particu-larly low CO2 abundances. This, combined withincomplete hydration, suggests that 504B crusthas not experienced all of the carbon uptake itwill see throughout its lifetime, and for that reasonit is not a reliable reference point for extrapolatingto global mass balances. This is acknowledgedin the global carbon budget of Alt and Teagle(1999) who ignored the data from the upperportion of 504B.

3.15.4.2 Chemical Fluxes between Oceanic Crustand Seawater: Methods andUncertainties

Geochemical exchange between seawater andthe oceanic crust varies substantially as a functionof depth and lithology, and for this reasonalteration processes are considered separately fordifferent portions of the crust in the 7.1 km ocean-crust reference section in Figure 2. Chemicalfluxes are given as D values, the differencein measured and original chemical inventory ing/100 g fluxes (at constant Ti), independently foreach depth interval in the oceanic crust. Anaverage altered MORB may be obtained by addingthis flux to the fresh MORB listed in Table 2.

The flux estimates in Table 2 primarily derivefrom the data from sites 417D and 418A for theupper crust, and site 735B for gabbros of the deepcrust, because these are the only drill sites withhigh recovery rates and total depths greater than300 m that sample crust older than 10 Ma, whichis considered a minimum age for “mature” ornear-mature ocean crust alteration. Study of siteswith a critical depth of penetration (.100 m?) ismore likely to help us understand their truebehavior than studies of very shallow siteswhere it is difficult to establish the overallalteration behavior. In particular, sites 417A,417D, and 418A allow for an in-depthunderstanding of localized anomalies.

For sites 417/418, the fluxes are calculated fromthe “super” composite, which is dominantlycomposed of materials from sites 417 and 418Aand offer a grand average for these sites. In itsupper 200 m, this composite includes about 20%of the oxidatively weathered materials from site417A. Sites 417/418 and 735 are the only deepdrill holes displaying average recovery rates.70%, allowing a quantitative reconstruction of

the alteration inventory. The main sources of datafor these sites come from Staudigel et al., 1989,1995, 1996; Smith et al., 1995; Spivack andStaudigel, 1994; Hart et al., 1999; Bach et al.,2001. However, for the intermediate crustalsections, data from site 504B was also utilized(e.g., Alt et al., 1993a,b; Alt and Teagle, 1999;Chan et al., 2002). This procedure was adoptedbecause there are no other data from in situ crust.It is important to note that use of these data leadsto increased uncertainties as alterations may notbe complete and the recovery rates are very lowfor 504B. The results from these deep-sea coresare also compared to evidence from the Troodosand Samail ophiolites (Bednarz and Schmincke,1989, 1994; Bednarz, 1989; Bickle and Teagle,1992; Spooner et al., 1977). However, in manycases, fluxes were extrapolated between site 735and Sites 417/418, using ophiolite data or datafrom 504B only for comparison. This procedurewas necessary for a variety of reasons includinglack of data, incomplete alteration due to youngoceanic-crustal age, and to produce the mostconservative estimate.

Obtaining realistic errors is one of the mostdifficult, yet most crucial problems in all fluxestimates. Such errors can be approximatedthrough an independent error analysis for severalfactors that are involved in estimating fresh andaltered rock composition. There are uncertaintiesarising from petrographic observations, in thechoices of representative samples, recovery ratebiases, and analytical errors. In most casesanalytical errors are a relatively minor source ofuncertainty, and they are typically rather welldocumented. Probably the most crucial analyticaluncertainty is in acurately determining the tita-nium concentration that is used as a normalizingfactor to account for open-system behavior. Thisuncertainty directly relates to an error in thefluxes, and thus fluxes are difficult to constrain tobetter than ,1% of the whole rock abundance of aparticular element.

Uncertainties in phenocryst abundance esti-mates are a significant source of error, particularlyfor major and highly compatible elements in themain phenocryst phases. In most cases related toocean drilling samples, phenocryst abundances ina particular sample are rarely known to better thanabout 5 vol.%. The most abundant phenocrysts areolivine and plagioclase, which contain theelements magnesium, silicon, calcium, aluminum,and nickel.

Detailed uncertainty estimates for majorelements are given in Staudigel et al. (1996,table 6). An indication of the overall reprodu-cibility of flux estimates can be obtained bycomparing independent data from different siteshaving similar alteration behavior, like site 417Dand 418A. The H2O and CO2 contents for

Chemical Changes in Altered Crust Composition Due to Hydrothermal Processes 523

composites at sites 417D and 418A as a functionof depth are compared to the large scale, bulkcomposition at these sites in Figure 4. CO2

abundances at these two sites track each othervery well, with an uncertainty between these sitesof ,0.5%. In contrast, the two H2O curves varymore widely (^1%). Close agreement betweencomposites from these sites suggest a high degreeof reproducibility, while a large scatter suggestssubstantial variation between sites or a problemwith reproducibility. In the following analysis thisscatter within and between sites is used to indicatehow robust these flux estimates are for the upperoceanic crust.

3.15.4.3 Chemical Fluxes

H2O. When discussing the water contents ofthe oceanic crust it is important to know that itmay be found in three forms that are measureddifferently and that are relevant for differentapplications. Ocean crust contains (i) formationwaters filling pore spaces, (ii) exchangeablewater in minerals like clays or zeolites, and(iii) water bound to crystalline structures ordissolved in silicate glass. Formation water canbe estimated from the pore space available in theformation, which is typically done by boreholelogging. Exchangeable water (H2O2) is typicallydefined as the water expelled at 110 8C, eventhough the release of this type of water is rathercontinuous and some fraction may occur attemperatures higher than this. Crystal-boundwater is referred to as H2Oþ and it is determinedin rock powders by titration or chromatographyfollowing ignition at temperatures .800 8C, or,in situ by spectroscopic methods. In mostgeochemical studies H2O is understood to beH2Oþ, in particular when H2O2 is not reported.Non-degassed, fresh MORB has approximately0.2 wt.% H2O. The amount of formation water inthe upper crust probably scales directly with theporosity of the crust, as it can be determined byborehole logging. The upper crust has about10 vol.% porosity (e.g., 13% in site 417D;Salisbury et al., 1979), which would yieldroughly 4 – 5 wt.% formation water for theupper crust. There are insufficient numbers ofH2O2 measurements to estimate the looselybound water in these sections.

Staudigel et al. (1995) estimated a H2Oþ flux forthe upper crust of about 2.81 g/100 g. It isinteresting that the average H2Oþ of compositesfrom the very top of cores 417D and 418A havevalues that are significantly below the values at100–200 m depth, and that site 418A has valuesthat are substantially above the values of 417D,giving an approximate uncertainty of ,0.5 wt.%for the supercomposite. Taking this average H2Oþ

as typical of the upper crust and including theestimate of formation water given above yields atotal of 7–8 wt.% H2O. Assuming ,2–3 wt.% forH2O2 suggests that the upper oceanic crust storesclose to 10 wt.% water, a water content that is notmuch different from those of pelagic sediments.This value is likely to decrease substantially withdepth, because pore space decreases with depthand because higher-temperature metamorphicassemblages contain less water.

An estimate of water content of the gabbrosection can be made using data from site 735B.Bach et al. (2001) suggest that gabbros from site735B took up about 0.11 g H2Oþ/100 g of rock.Formation water can be constrained from poro-sities in this site. Core materials from site 735Bhave porosities of 0.7 vol.%, and in situ porosity islikely to be twice as high, probably about1.5 vol.% (Dick et al., 2000), which is substan-tially less than that of the upper crust. Filling1.5 vol.% of the rock with water yields about0.5 wt.% of H2O, suggesting that the gabbros arelikely to store slightly less than 1 wt. % water—anorder of magnitude less than that present in theupper crust. The mass balance uses the H2Oþ fluxof Bach et al. (2001), which should be close to atotal average for the lower crust (even though735B penetrated only through its upper portion).

Water contents at intermediate depths areextrapolated from the top of the crust downwardusing the internal variation in sites 417/418 andthe data for minimal alteration from site 504B andsite 735B. The H2Oþ fluxes into the lowerextrusives and sheeted dikes are estimated at2.10 g/100 g and 1.5 g H2Oþ/100 g, respectively(see Figure 4(a)). Adding up and averaging theseindividual estimates, the oceanic crust as a wholetakes up ,0.45 g H2Oþ/100 g of rock. The totalwater inventory of the whole crust (includingH2Oþ, H2O2, and formation water) is derivedsimply by scaling (in a similar fashion as in theupper crust), at ,1.5 wt.%.

CO2. Degassed MORB has ,0.12 wt.% CO2,whereas undegassed MORB has ,0.45 wt.% CO2

(Table 2; Gerlach, 1989). The higher CO2 inaltered MORB is due to CO2 addition to the crustfrom seafloor alteration subsequent to initialoutgassing. This alteration-related addition ofCO2 to the oceanic crust brings its total inventoryup to levels exceeding the original non-outgassedcarbon contents (Staudigel et al., 1989). Most ofthis hydrothermal CO2 inventory in the oceaniccrust is added in the form of carbonates,particularly in the upper 600 m (Figure 3(b),Staudigel et al., 1989). In the estimate of carbonfluxes (Table 2), data for the upper 600 m of sites417/418 yielded a CO2 uptake of 3.26 g/100 g(Staudigel et al., 1989). Based on the scatter ofthe depth composites and the discrepanciesbetween sites 417D and 418A, the error on these


estimates is roughly 0.3 g/100 g. This estimatefrom sites 417 and 418 was confirmed by Alt andTeagle (1999) using data from sites 843(90 m penetration) and 801(150 m penetration),which yielded 2.4 wt.% and 3.55 wt.% CO2.,respectively.

For the lower crust we use carbon uptakeestimated by Bach et al. (2001) for the gabbros atsite 735B, yielding a CO2 flux of 0.06 g/100 g intothe 1,500 m gabbro section. Lower extrusives andsheeted dikes from site 504B apparently have notseen the complete cycle of alteration, and there-fore the very low CO2 values from these sites arenot robust indicators of the total carbon budget ofthe intermediate crust. To estimate the carboninventory of the intermediate levels of the oceaniccrust, we interpolate between the extrusives andthe gabbros, which yields an estimate of 0.5 gCO2/100 g for the lower extrusive section, slightlyabove the value of Alt and Teagle’s (1999)transition zone of site 504B. For the sheeteddikes we use 0.06 g/100 g, the value of the gabbrosection, as a conservative minimum estimate.

The total uptake of the crust is calculated here as0.355 g/100 g, which is slightly above the estimatemade by Staudigel et al. (1989), and also similar tothe total carbon uptake of the oceanic crustinferred by Alt and Teagle (1999) and for theTroodos ophiolite (Bednarz and Schmincke,1989). The above extrapolation of data forintermediate depths (lower extrusive crust, sheeteddikes) does not contribute major uncertainties tothe total flux estimate, because most of the carboninventory is located in the upper 600 m of thecrust, and, therefore, most of the uncertainties liein this depth interval.

Oxygen isotopes. d18O in ocean-crust studies istypically defined as the per mil deviation in18O/16O ratio of a rock relative to a standard meanocean water (d18OSMOW) and it is widely used tounderstand ocean-crust alteration processes. FreshMORB has an d18OSMOW value of þ5.7‰,and water – rock interaction with seawater(d18OSMOW ¼ 0‰) at low temperatures increasesthe value, while high-temperature alterationdecreases it. Muehlenbachs and Clayton (1972)drew attention to this relationship and suggestedthat hydrothermal alteration of the crust maybuffer the oxygen isotopic composition ofseawater. Oxygen is the major component in theoceanic crust, and therefore, changes in d18O are arather profound indicator of hydrothermalalteration.

The silicate portion of the upper oceaniccrust at sites 417/418 has a d18O of ,þ10‰(þ9.98‰, Staudigel et al., 1995), substantiallyelevated from the original magmatic value ofþ5.7‰. The values for the composite crustdecrease with depth from ,þ11.5‰ in the upper100 m, down to ,þ8.5‰ between 400 m and

600 m depth, roughly defining an uncertaintyband of ,^1d18O units. This uncertainty is rathersmall when compared to data from other ocean-crust sections. For example, the same depthinterval in the Troodos and Samail ophiolitesdisplays values ranging from ,8‰ to 14‰ (e.g.,Spooner et al., 1974; Gregory and Taylor, 1981;Stakes and Taylor, 1992), which is substantiallylarger than the 417/418 composite range, butthese ophiolites also display a general decrease ofd18O with depth. Site 735B shows a ratherinteresting variation, even though the data aremore scattered. Its uppermost 500 m display arange of 3–7‰, with an average d18O of 4.35‰(Hart et al., 1999), well below the value of freshMORB, and clearly indicating relatively highreaction temperatures. Deeper sections of 735Bdisplay heavier d18O values, also with a signifi-cant scatter, but with a higher average value, onthe order of 5.8‰. This is not very different fromthe original magmatic value (Bach et al., 2001).This pattern of very low d18O values at the top ofthe gabbros that slightly increase with depth isquite similar to the variation seen in the Troodosand Samail ophiolites (Spooner et al., 1974;Gregory and Taylor, 1981; Stakes and Taylor,1992).

The d18O of the average crustal section inTable 3 derives from the average d18O for the top600 m of sites 417/418 (d18O ¼ 10‰), with d18Odecreasing to 8‰ in the lower extrusives, 4‰ inthe sheeted dikes and then increasing again to5.8‰ in the gabbros from site 735B. This overallvariation is solely derived from drill holes innormal oceanic crust, but it is remarkably similarto the patterns observed in ophiolites. Thissuggests that 735B is actually quite an appropriateanalogue for hydrothermal alteration in “true,”deeply buried gabbro sections in normal oceaniccrust. It appears that most hydrothermal systemspenetrate into the top few hundred meters ofgabbros, whether or not the section carries anyoverlying extrusives and dikes.

Strontium. The 87Sr/86Sr ratio is also a verypowerful indicator of hydrothermal alteration inthe oceanic crust, because fresh crust is unradio-genic (87Sr/86Sr ¼ 0.7025–0.703; Chapter 3.13)while seawater is radiogenic (0.7092; Chapter6.02). Seawater 87Sr/86Sr also changes with time.One-hundred-million year old seawater was lessradiogenic (0.70735) than today’s, and 87Sr/86Srhas been continuously increasing since then. Theisotopic composition of reaction products quanti-tatively reflects the strontium contributions fromseawater and from basalt. The reference sites 417/418 and 735B both display primary strontiumratios close to 0.7029 (Staudigel et al., 1981; Hartet al., 1999; Bach et al., 2001). The average87Sr/86Sr ratio at sites 417/418 is 0.704575, with arather systematic down-hole variation that closely


mirrors the behavior of H2O. Alteration-relatedincreases in 87Sr/86Sr at the top of the crust arerelatively modest, become more pronounced at adepth of ,200 m, and then decrease with depthagain, to its lowest values at 550 m (Staudigel et al.,1995). Hart et al. (1999) and Bach et al. (2001)showed that site 735B has generally very low87Sr/86Sr (0.70295), with some small-scale varia-tion that might resemble several cycles of decreas-ing 87Sr/86Sr with depth. 87Sr/86Sr values from sites417/418 are high relative to those from site 504B(e.g., Barret and Friedrichsen, 1982; Friedrichsen,1984; Kawahata et al., 1987), but are relatively lowwhen compared with 87Sr/86Sr of the Troodosophiolite (Bickle and Teagle, 1992). The formermight be due to incomplete alteration at 504B andthe latter due to higher initial isotopic ratios in freshbasalts in the Troodos ophiolite. The trends in87Sr/86Sr at sites 417/418 and in the Troodosophiolite are used to extrapolate the 87Sr/86Sr of thelower extrusives (87Sr/86Sr ¼ 0.704) and thesheeted dikes (0.7035). For the gabbros, 87Sr/86Sris assumed to be 0.70295, which is only slightlyelevated from the original values (Hart et al., 1999;Bach et al., 2001).

Strontium isotope ratios and abundances insamples from the oceanic crust may be used todetermine the complete chemical mass balance ofstrontium exchange between seawater and basalt,including the loss of basaltic strontium tohydrothermal solutions, and uptake of basaltic orseawater strontium from hydrothermal solutions.A mass balance of this exchange can be made infour steps. (i) The relative amount of basaltic andseawater strontium in altered basalt can bedetermined from the measured 87Sr/86Sr of analtered sample as a (linear) mixture of strontiumfrom the two end members, the (contempora-neous) seawater and basalt. (ii) The inventory ofthe basaltic and seawater strontium in an alteredsample (in mg/kg) may then be determined fromthe above ratio of seawater and basalt in thesample and the total strontium abundancemeasured for this sample. (iii) Seawater strontiumaddition to the basalt is given directly by theseawater strontium inventory calculated inStep (ii). (iv) The determination of flux of basalticstrontium in or out of an altered sample is morecomplicated because it has to be related to theoriginal inventory of strontium. It is determined asthe difference between the original basalticinventory and the basaltic strontium present inthe altered sample.

Using this type of mass balance, it can be shownthat on average, basalts from the upper crust atsites 417 and 418 have lost 10 mg/kg of basaltstrontium to seawater and gained 32 mg/kg ofseawater strontium, resulting in a total flux of22 mg/kg of strontium into the basalt (Staudigelet al., 1996). Gabbros at 735B show very large

changes in strontium abundance between freshand altered sample pairs, but these changes showboth gains and losses of strontium by up to 25% ofthe total inventory (Bach et al., 2001). Hart et al.(1999) estimated from strip composite samplesthat the upper portion of 735B shows no or verylittle change in strontium abundances. In Table 2,strontium fluxes are derived from the data of sites417/418 and 735B for the upper and lower crust,respectively, effectively tapering the rather highstrontium fluxes into the upper crust to the nearzero strontium fluxes in the gabbros. A slightstrontium loss is assumed for the gabbros, becausethe breakdown of calcic plagioclase is likely toliberate substantial quantities of strontium, someof which may be removed by hydrothermalsolutions. Over the total crust, these elementalchanges add up to a slight uptake of strontium, of1.4 mg/kg of rock, which is extremely small whencompared to the overall fluxes of strontium intoand out of the crust. Thus, seafloor alteration cansubstantially influence the isotopic composition ofseawater, but not significantly change the stron-tium inventories in either reservoir (see Palmerand Edmond, 1989; Elderfield et al., 1999).

Sulfur. Sulfur is a key element in hydrother-mal processes. Seawater introduces sulfate intothe hydrothermal system that precipitates anhy-drite during heating (e.g., Shanks et al., 1981).Basalt is very reduced and contains sulfurdissolved in melt and immiscible sulfur metalglobules (Mathez, 1976, 1980). The reaction ofseawater with basalt results in the precipitationof sulfides, mostly pyrite, as one of the mostcommon reaction products. There are only veryfew systematic sulfur and isotope studies of theoceanic crust, in particular those by Alt (1995)on ODP site 504B, Bach et al. (2001) on site735B, and Alt (1994) on the Troodos ophiolite.The pre-alteration sulfur abundance at site 504Bis 960 ppm. Overall, basalts appear to be asource of sulfur to hydrothermal solutions,delivering ,0.1–2 g of sulfur per 100 g ofbasalt (recalculated from Alt, 1995) and sulfur isprecipitated in veins (sulfide and anhydrite), inparticular in mineralized zones at the bottom ofthe extrusives and top of the dikes (Alt, 1995).Alt (1995) determined the volumes of veinsin 504B and estimated a total flux of0.15 £ 1012 g yr21 from the oceanic crust, cau-tioning that this is the only site where such amass balance has been done and citing relativelylarge errors. However, the total sulfur flux inaltered oceanic crust is relatively small whencompared with the sulfur in high-temperature

hydrothermal fluxes ((10.3–25.7) £ 1012 g yr21;Von Damm et al., 1985; Sleep, 1991) and riverfluxes (28.5 £ 1012 g yr21; Holser et al., 1988).Alt (1995) suggests that the discrepancy in thehydrothermal fluxes is probably best explained


by anhydrite precipitation during black-smokeractivity, and subsequent dissolution from low-temperature fluids. In Table 2 a sulfur lossof 0.06 g/100 g (Alt, 1995) is estimated for theextrusive section and a sulfur gain of 0.02is estimated for the gabbros at site 735B(Bach et al., (2001). The total sulfur flux is0.0037 g/100 g from basalt to seawater.

The d34S of fresh MORB isþ0.1‰ (Sakai et al.,1984). The calculated d34S of 21.9‰ for theextrusives, þ1.3‰ for the dikes (anhydrite veinsnot included) and þ0.13‰ for the gabbros(Table 2) is taken from Alt (1995), who estimatesthe average d34S for the total oceanic crust asþ0.9‰ per mil.

Lithium. Lithium is a very sensitive indicator ofhydrothermal alteration processes, because it isadded during low-temperature alteration, mostlyin smectites, and is leached during high-temperature alteration (e.g., Seyfried et al.,1984; Von Damm et al., 1985). In addition,lithium isotopes reflect seawater additions to theupper crust, much like strontium and oxygenisotopes do. The most comprehensive study oflithium abundances and isotopic ratios wasrecently presented by Chan et al. (2002) for site504B. Considering the low recovery rates at thissite, this study should be considered to providea minimum baseline of lithium isotopic altera-tion, with an emphasis on the early portionof alteration. Unaltered MORB has anaverage lithium abundance of 4.5 ^ 1.5 ppm(Chan et al., 1992; Ryan and Langmuir, 1987).The least altered rock from site 504B has ,3 ppmlithium, the average lithium inventory in thevolcanics is ,5.8 ppm, indicating an uptake of2.8 mg/kg. The average lithium content of thesheeted dike (þ the transition zone) is ,2.5 ppm,indicating a loss of 0.5 mg/kg. Chan et al. (2002)suggest that the initial lithium content of site 735is ,3 ppm and the measured value is 1.0,suggesting a loss of 2 mg Li/kg. In total, thecrust loses ,1.09 mg Li/kg of basalt.

d 7Li, the per mil deviation of the 7Li/6Li ratiofrom the NBS SRM L-SVEC, of fresh MORB isþ3.4 – 4.7‰ (Chan et al., 1992; Ryan andLangmuir, 1987) and rises steeply with the degreeof low-temperature alteration (Chan et al., 1992),reflecting uptake of heavy seawater lithium(þ32‰). d 7Li decreases with depth in site504B, from an average of ,6‰ in the extrusivesto ,2‰ in the sheeted dikes. This reflects additionof seawater lithium to the upper crust and removaland fractionation of lithium in the deeper crust(Chan et al., 2002).

Boron. Boron concentrations in fresh MORBare ,0.5 ppm (Spivack and Edmond, 1987; Ryanand Langmuir, 1987) and 26.2 ppm in averageupper crust at sites 417 and 418 (Smith et al.,1995). Boron concentrations in altered crust show

a significant decrease with depth, from ,50 ppmin the upper 100 m of 417/418 to 8.5 ppm at 500 m(Smith et al., 1995). Subtracting the initial basalticboron inventory from the altered extrusiveinventory yields a boron uptake of 25.7 mg/kgfor the upper 600 m of extrusives. Extrapolatingthe 417/418 data downward, the lower extrusiveshave ,4.5 ppm boron, and an uptake of 4 mg/kg(Table 2). An uptake of 5.6 mg/kg for the sheeteddikes and 2.2 mg/kg for the gabbros is estimatedusing Layer 2B and Layer 3 estimates for boronfrom Smith et al. (1995; 5.9 ppm and 2.7 ppm,respectively). The total uptake is thus 4.81 mg/kgfor the whole oceanic crust. d11B (i.e., the per mildeviation of the 11B/10B ratio from the NBS SRM951 standard) in the upper crust can increasesubstantially during seafloor alteration. Hyalo-clastites may reach d11B values up to þ5.4‰, butthe average crust is only slightly elevated(þ0.8‰) relative to the least altered compositesof flows (0.5‰, Smith et al., 1995). Gabbros fromthe upper portion of site 735B have markedly

elevated d11B of 7.35‰ (Hart et al., 1999). d11Bfor intermediate depths are extrapolated fromthese values to be 1‰ for the lower extrusives and2‰ for the sheeted dikes, resulting in an overallvalue of d11B ¼ 5.6‰ for the total crustal section.

Potassium, rubidium, and caesium. They areamongst the most sensitive indicators of water–rock exchange. Like their sister alkali element,lithium, these elements are taken up by basaltsduring low-temperature alteration reactions andleached from basalts during high-temperaturereactions. The upper 600 m of crust at sites 417and 418 records gains of 0.54 g/100 g K2O,10.3 mg/kg rubidium and 0.183 mg/kg caesium(Staudigel et al., 1995, 1996; Hart and Staudigel,1989). Surprisingly, the gabbros at site 735B allshow gains of these alkalis as well: 0.01 g/100 gK2O, 0.7 mg/kg rubidium and 003 mg/kg caesium(Bach et al., 2001). This may reflect the anomalousnature of the crust at 735B, because it is directlyexposed to seawater, or it may be a “typical” late-stage low-temperature alteration overprint thataffects all of the oceanic crust, even if it waspreviously depleted in potassium, rubidium, andcaesium from high-temperature reactions at theridge axis. For this reason it is not clear whether thelower oceanic crust displays slight gains or near-complete losses in these elements. Given the ratherlow initial inventories of potassium, rubidium, andcaesium in oceanic gabbros and the very largegains in the upper crust, the above uncertainties inthe gabbro inventories have a relatively smalleffect on the alteration budget of these elements forthe whole crust. For these reasons, the data fromsites 417/418 are used for the upper oceanic crustand depth trends at these sites are used toextrapolate the values for the lower 400 m of theextrusives. The sheeted dikes and the gabbros are


proposed to have experienced no net change inpotassium, rubidium, and caesium. Using averagevalues of 0.05 g/100 g K2O, 1 g/kg rubidium, and0.03 g/kg for caesium results in total fluxes of0.0485 g/100 g of K2O uptake, 0.927 mg/kg ofrubidium uptake, and 0.0172 mg/kg of caesiumuptake for the whole crust.

Uranium. Uranium content is also enhanced inthe upper oceanic crust during alteration. Theuranium content of the altered upper crust is0.30 ppm (Hart and Staudigel, 1989; Staudigelet al., 1996). Uranium is slightly depleted in theupper part of site 735B (Hart et al., 1999), butslightly enriched overall in 735B (Bach et al.,2001). Data for the top 600 m of site 417/418 aretherefore used for the upper crust, and the grandaverage of 735B is used for the lower crust. As inthe previous estimates, the internal variation in417/418 and 735B are used to extrapolate theintermediate crust, yielding a total uptake of0.037 mg/kg of uranium in the crust.

Rhenium and osmium. There are no data onbehavior of Re/Os during alteration of MORB, butsome work has been done on pillow lavas fromthe submarine seamount sequence in thebasement complex of La Palma (Canary Islands;Marcantonio et al., 1995). There, altered pillowrinds have 187Os/188Os ratios of 0.40–0.43, whilethe pillow interiors generally have 187Os/188Osvalues between 0.14–0.19. Schiano et al. (1997)found highly variable 187Os/188Os ratios in mid-ocean ridge glasses and cannot rule out potentialassimilation of manganese crusts with high187Os/188Os. Gabbros from site 735B suggestisotopic variations consistent with these changes(Hart et al., 1999). Thus, seafloor alteration islikely to increase the 187Os/188Os significantly insome rocks, in particular in the upper crust(Marcantonio et al., 1995), but it is too early todetermine whether this effect is of any more thanlocalized importance.

SiO2 is amongst the most difficult element totrack during alteration of the oceanic crustbecause it shows major variations due to igneousprocesses and because the majority of alterationreactions involve silicates. The upper extrusivecrust at 417/418 gains about 1.2 g/100 g SiO2. Theuncertainties are expectedly very large, including^0.18 g/100 g in the estimate of the fresh rockcomposition, and a variance of nearly 4 g/100 gin fluxes at various depths of sites 417D and 418A(Staudigel et al., 1995, 1996). Bach et al. (2001)estimated that gabbros of site 735B gained about0.1 g SiO2/100 g gabbro, and the scatter in theirdata suggest that this value is also rather uncertain.We extrapolate between these values and estimatethe fluxes to be 0.5 g/100 g for the lower extrusivecrust and 0.25 g/100 g for the dikes (Table 2).However, these values are likely to have errors onthe order of ^2 g/100 g, which are extremely large

when compared to fluxes at hydrothermal ventsand in the ocean in general.

Al2O3. Fresh MORB contains, on average,15.26 wt.% Al2O3 (Hofmann, 1988; see alsoChapter 3.13). Alteration of the upper crust causesa gain of about 1.94 g/100 g of Al2O3 (Staudigelet al., 1996), based on individual fluxes at 417Dand 418A that vary from 21 g/100 g to5.0 g/100 g. This range indicates with significantuncertainty, partly due to the large uncertainty inthe fresh estimate (^0.37 wt.%, mostly fromuncertainties in plagioclase phenocryst distri-bution) and to local redistribution of aluminumwithin the holes. An analytical uncertainty of ,1%in the estimate of the altered rock compositionwould contribute an error of ^0.15 g/100 g to thefluxes. The lower crust Al2O3 content is unlikelyto be affected by alteration.

FeOtot in unaltered crust is 10.43 wt.%(Hofmann, 1988; see also Chapter 3.13). Theupper crust at 417 and 418 loses about 0.01 g/100 g; individual composites may show losses to0.85 g/100 g or gains of up to 1.01 g/100 g of up(Staudigel et al., 1996). These variations are mostlikely to represent uncertainties in the estimatesrather than true fluxes. Iron concentrations inMORB are very high and variable while seawatercontains very little iron. The analytical uncertaintyin the determination of iron alone is on the order of0.1 wt.%, which is almost an order of magnitudehigher than the flux. This shows that iron isessentially immobile and that the fluxes are toolow to detect beyond the uncertainties of crustalmass balances. However, due to the extremely low-iron inventories in seawater, even the smallestfluxes would be quite significant for the seawaterbudget, and, thus, the oceanic crust may deliversubstantial fluxes to seawater. Given the concen-tration gradient, these fluxes are likely to move ironfrom the crust to the oceans but there is noconvincing data from the oceanic crust that canprove this.

MnO shows no change in the upper extrusives,but fluxes for individual composites range from0.04 g/100 g to 0.05 g/100 g (Staudigel et al.,1996) suggesting very large uncertainties that donot allow us to estimate meaningful fluxesindependently. The very low manganese concen-tration in seawater and the enrichment of manga-nese in hydrothermal waters suggests that it islikely that manganese leaves the crust, eventhough local enrichments of manganese are likelynear the top of the extrusive oceanic crust due todeposition of manganese crusts and Fe–Mnhydrothermal deposits.

MgO is a key element in the discussion ofhydrothermal systems. It is an important elementin many igneous phases and in alteration phases.The depletion of magnesium in hydrothermal ventfluids was originally seen as evidence that the


oceanic crust is a major sink for magnesium fromriverine input to the oceans (Edmond et al., 1979;Van Damm et al., 1985). However, flux estimatesfrom rock analyses do not to support thiscontention, at least not in a very consistentmanner. Data from sites 417/418 and 735Bsuggest that, on average, rocks lose moremagnesium than they take up (0.2 g/100 g inboth cases, Staudigel et al., 1996; Bach et al.,2001). Dredged samples of abyssal peridotitesuggest that they lose up to 5 wt.% magnesiumdue to alteration (Snow and Dick, 1995). Mag-nesium uptake was reported at the Troodosophiolite (Bednarz, 1989; Bednarz andSchmincke, 1989), and for site 504B (Alt et al.,1996). Part of this discrepancy may be due to thevery large uncertainties in these fluxes. At sites417/418, fluxes of individual depth compositesvary from 21.1 g/100 g to þ2.01 g/100 g, and theuncertainty in the fresh rock estimate is 0.19 g/100 g. Thus, the data from 417/418 actuallypermit a small gain in MgO, but the uncertaintiesare large. The estimate in Table 2 uses the 417/418data as representative of the upper oceanic crustand makes the most conservative estimate oflower crust/sheeted dikes of a magnesium flux ofzero (Table 2). The crustal budget of magnesiumis similarly as poorly constrained as the budget foriron, but it is likewise clear that there really are novery large fluxes. A major difference betweenmagnesium and iron is that magnesium has amuch larger inventory in seawater, which requiresmuch larger fluxes to make it significant for theseawater–ocean-crust budget. These large fluxesare not observed.

CaO. Calcium is also an element that is veryimportant to global geochemical cycles, as theoceanic crust is considered to be an importantsource of Ca2þ (e.g., Berner et al., 1983). Thereare several competing processes that control thecalcium content in the upper crust. Calcium is lostduring glass alteration (Staudigel and Hart, 1983)and calcic feldspar breakdown, and is taken upduring calcium carbonate precipitation. Sites417/418 display an uptake of 2.1 g/100 g^ 1 g/100 g CaO (Staudigel et al., 1996), with an overalldecrease with depth. Alt et al. (1996) report asmall loss of CaO at site 504B, which is consistentwith the limited precipitation of carbonate there,while much of the calcic feldspar is replaced.Extrusives in the Troodos ophiolite drill coresconsistently show loss of CaO: 8.08 g/100 g in theupper crust and 3.34 g/100 g in the lower crust(Bednarz, 1989). The flux values for CaO inTable 2 use site 417/418 data for the top of thecrust, and site 735B data for the gabbros. Thesesites are considered the most representative of theoceanic crust. Intermediate values in the lowerextrusives and sheeted dikes are extrapolated fromthese data and tapered to zero fluxes, like CO2,

largely because much of the calcium depositionappears to be linked to carbonate precipitation.This yields an uptake of 0.03 g/100 g in the lowerextrusives and an uptake of 0.04 g/100 g in thesheeted dikes. Altogether these calcualtions sug-gest a slight positive flux of calcium into the crustof 0.134 g/100 g. Most evidence suggests thatcalcium contents do not show much net change,even though it is quite likely that the exchange ofcalcium between seawater and basalt is very high,probably quite similar to the behavior of strontium.

Na2O is a mobile element that is involved inimportant alteration reactions in the shallow crust(loss during glass alteration, gain from formationof zeolite), in the deeper crust (formation ofzeolites, albite, and halite) and has a substantialinventory in seawater, which is at times enhancedby boiling of water in high-temperature hydrother-mal systems followed by the precipitation of halite.The oceanic crust appears to gain small quantitiesof sodium, on average, particularly in the uppercrust (0.15 g/100 g in 417/418; Staudigel et al.,1996). Bach et al. (2001) suggest a gain of 0.03 g/100 g for the gabbros at site 735B. A gain of 0.15 g/100 g for the lower extrusives and the dikes is usedin Table 2. This is intermediate between the smalleruptake at site 504B (0.09 g/100 g; Alt et al., 1996)and the higher uptake at the Troodos ophiolite(1.9 g/100 g lower extrusives; 0.421 g/100 gsheeted dikes; Bednarz, 1989). Like other majorelements, however, it is obvious that the gains orlosses of sodium in basalts are quite variable anduncertainties are high.

A large number of elements are immobileduring alteration of the oceanic crust. These aregenerally elements that are insoluble in seawaterand include titanium, which is commonly used asan immobile reference element to constrain over-all losses or additions in a chemically open systemand REEs, hafnium, niobium, zirconium, andthorium. Similarly, the neodymium and hafniumisotopic composition of oceanic crust is rarelyaffected by alteration. However, significant mobi-lity of otherwise immobile elements has beenidentified in extremely altered basalts dredgedfrom some seamounts (e.g., Cheng et al., 1987).Staudigel et al. (1996) observed distinct but minorchanges in 143Nd/144Nd and in Ce/Cep in the mostaltered composite samples from sites 417 and 418,consistent with the addition of REE in very largequantities of seawater or moderate amounts ofsediment particulates (Staudigel et al., 1996).

3.15.5 DISCUSSION

3.15.5.1 Hydrothermal Fluxes: Rock Dataversus Fluid Data

The estimates of elemental changes due tohydrothermal alteration of oceanic crust in Table 2

Discussion 529

can be re-cast into global geochemical fluxesusing the ocean-crust production rate discussedabove (see Figure 2). These fluxes can then becompared with river fluxes and two types ofhydrothermal fluxes: axial hydrothermal fluxes orflank fluxes following Wheat and Mottl (2000).

3.15.5.1.1 Uncertainties

Errors in elemental fluxes derived from crustalestimates are larger than, or similar to, the value ofthe actual flux estimate for many of the majorelements: silicon, aluminum, iron, manganese,magnesium, and sodium. The fluxes for theseelements are thus poorly constrained, but theseestimates do serve as conservative bounds onthe fluxes. Unfortunately, these bounds overlapthe fluxes derived from hydrothermal fluid dataand river data. For this reason, current ocean-crustflux estimates do not provide independent evi-dence for the magnitude of hydrothermal fluxesin the geochemical cycle for these elements.Within the bounds of these uncertainties, the dataindicate that ocean floor hydrothermal processesmay balance (or compound!) missing globalfluxes of these elements. For these reasons, theseelements are not discussed here in any detail.

The flux estimates of CaO are slightlybetter known than the forgoing major elements,but the uncertainty of CaO flux in the upper oceaniccrust is still almost two-third of its value. Muchlarger uncertainties are typical of the deeperextrusives and sheeted dikes, where spilitizationis characterized by near-complete loss of calcium.All other element fluxes reported here haveuncertainties that are less than 30% of their value.

3.15.5.1.2 Bulk fluxes

A comparison can be made between fluxesderived from the rock record versus thoseobtained from hydrothermal fluids for elementsfor which well-constrained rock data exist. Theseelements are calcium, potassium, CO2, and sulfur(Table 2). This is a somewhat random collectionof elements and the comparisons thus derivedmay not be “typical” of all elemental fluxes, butthey do give some first-order comparisons. Thecrustal fluxes of these elements add up to3.6 £ 1014 g of seawater calcium, potassium,CO2, and sulfur per year, into the crust, whilethe sum of the flank and axis hydrothermal fluxesgive a net flux out of the oceanic crust rangingfrom 1.8 £ 1014 g yr21 to 2.92 £ 1014 g yr21,depending on whether one uses the high or lowestimates for axial fluxes. Overall, the rock recordyields fluxes into the crust of the same order ofmagnitude but opposite sign to the fluid data,illustrating a rather fundamental disconnectbetween hydrothermal fluid data and ocean crust

alteration data. Ideally, fluxes derived from dif-ferent and complementary data sets should be thesame. What are the reasons for this discrepancy?Several contributing factors may be considered:

(i) The two methods sample different types ofprocesses. When Edmond et al. (1979) publishedthe first “global” geochemical fluxes based onblack-smoker data, Hart and Staudigel (1982)pointed out that there are major discrepancies withthe fluxes of rubidium and caesium from ocean-crust alteration data that appear to be largelycontrolled by low-temperature alteration. Fluxesbased on hydrothermal fluids are biasedtowards high-temperature processes, which mayunderestimate the total flux.

(ii) The two methods sample processes occur-ring at very different times in the hydrothermalhistory of the oceanic crust. The oldest hydrother-mal vent samples analyzed to date come fromBaby Bare seamount near the Juan de Fuca Ridge,on 3.5 Ma crust (Wheat and Mottl, 2000). At3.5 Ma, the oceanic crust has lost only half of itstotal convective heat loss (C. A. Stein and S. Stein,1994b). Hydrothermal fluxes from the rock recordat sites 417 and 418, for example, offer a completeset of chemical changes occurring during theentire hydrothermal history of the oceanic crust.

(iii) Oceanic crust has a substantial amount ofpore space produced during its initial emplace-ment, and new pore space is generated as the crustcools. All these pores are filled with seawater andsome additional water is taken up throughhydration of ocean-crust materials. The chemicalinventory of these pore waters are transferred intothe crust, producing a one-way flux of water intothe crust that is not accounted for through themeasurement of hydrothermal fluids. This fluxrepresents a baseline flux that should be subtractedfrom hydrothermal fluxes in order to calculatenet fluxes out of the crust. It was estimated abovethat oceanic crust may take up to ,1.5 wt.%water. This translates into addition of ,2 £1012 g MgO yr21, 0.5 £ 1012 g CaO yr21,1.5 £ 1013 g Na2O yr21, 1 £ 108 g Rb yr21,0.3 £ 108 g CO2 yr21, and 1 £ 1013 g S yr21.Overall, these fluxes are relatively small whencompared to the fluxes considered in this massbalance and are thus not likely to explainthe discrepancy between hydrothermal- andocean-crust-derived fluxes.

(iv) Another issue in reconciling fluid versuscrust fluxes relates to the nature of fluid circulationin the oceanic crust and the uncertainty in fluidpathways. With the exception of pore waters,fluids sampled from hydrothermal systems reflectrelatively high water volumes, and high water–rock ratios. Most of the crust however, is altered atlow water–rock ratios, in particular deep withinthe crust (.100m). Thus, it is possible that thealteration reflected in hydrothermal fluid data only


represents hydrothermal activity occurring in highwater–rock ratio reaction zones and in high water-throughput aquifers. By far the greatest volume ofoceanic crust is altered at much lower water–rockratios and by processes that may largely bediffusive. In addition, drilling recovery rates areparticularly low in regions that are most porous,and thus high water–rock ratio regions, which aretypically very porous, may be under-representedin crustal studies

(v) The role of sediments in mass balances ofseawater-ocean-crust chemical exchange remainslargely unexplored. Pore-water studies in sedi-ments on the ocean floor demonstrate that there ischemical exchange between sediments andthe basaltic oceanic crust (Lawrence and Gieskes,1981). The impact of these processes is notquantified but it is quite possible that thisexchange has a significant affect on the bulkcomposition of the oceanic crust. This is particu-larly important for sedimented ridges and laterstage alteration of the oceanic crust, when thecrust is typically covered by sediments. Chemicalexchange between sediment pore waters and theocean crust may have a profound impact on thechemical fluxes into the oceanic crust, but it is notincluded in the hydrothermal spring data.

At least four out of these five contributingfactors may substantially contribute to the dis-crepancies between fluid- and crust-derivedhydrothermal flux estimates.

3.15.5.1.3 Reconciling hydrothermal fluxesfrom fluid and rock data

The discrepancies discussed above suggest thatboth methods of flux determination may haveintrinsic problems. Flux data from hydrothermalvents are compromised for determining global fluxestimates because these fluids are derived fromyoung crust, mostly from deeper crustal levels.They selectively sample high water–rock ratioalteration processes and they ignore low-temperature fluid fluxes (,20 8C). Rock data areparticularly unreliable or not available forthe deeper crust and for high water – rockratio reactions. Low recovery rates, and largeprimary variability in many important elements(magnesium, silicon) produce significant uncer-tainties in the flux estimates for elements whosefluxes are particularly well determined in hydro-thermal fluids. The most important steps inreconciling these differences are to evaluate theweaknesses of both approaches and try to arrive atfluxes that use both methods in a complementaryfashion. Furthermore, a critical error analysis willhighlight the weaknesses of each of the methods,allowing the design of new experiments orapproaches that resolve these issues by reducing

a particular source of uncertainty, possiblyseparately for different element groups.

3.15.5.2 Impact of Ocean-crust Composition onArc Processes and Mantle Heterogeneity

Using average MORB or the range of compo-sitions of oceanic basalts (e.g., Hofmann, 1988;Chapter 3.13 and http://petdb.ldeo.columbia.eduLehnert et al., 2000), the fluxes derived here canbe applied to determine the average compositionsof oceanic crust that is subducted and recycledinto the mantle. These compositions thus influ-ence the composition of subduction zone magmas(see Chapter 3.18) and bear on the chemical massbalance of the mantle.

Whereas average fluxes are useful for definingthe global mass balance between mantle and theoceans, understanding the compositional diversityof subducted crust is important in constraining itsdehydration or partial melting processes duringrecycling (see Chapter 3.17). Such diversity isreflected in the more extensively altered compo-sitional domains in the oceanic crust such asvolcaniclastics, and the moderately altered flowcomposites of sites 417A, 417D, and 418B(Staudigel et al., 1995, 1996). Other extremecompositional domains include umbers, ophical-cites (calcite-basalt breccia), massive sulfides,epidosites, talcum deposits, or serpentinites, butthese are probably best studied in ophiolites.Particular mineralogical assemblages may showdistinct phase relationships during prograde meta-morphism. For example, ophicalcite or talcum hasbeen shown to display distinct phase relations thatis likely to control its dehydration or decarbona-tion behavior during subdution (Wyllie, 1978;Kerrick and Connelly, 2001). Volatile-rich oralkali-rich compositional domains may contributepreferentially to fluids that are extracted duringprograde metamorphism in subduction zones, orthey may melt prior to the surrounding lessalkalic, and volatile-poor rock. The latter scenariohas been considered in mantle melting models as aprocess that explains the isotopic variation inmantle-derived melts (Phipps-Morgan, 2001).MORB are generated by massive melting eventsthat mix melts derived from depleted mantle andrelatively enriched “plums” of recycled materials,while the smaller degrees of melting of oceanisland basalts are capable of extracting more ofthese heterogeneities, displaying the fuller extentof mantle heterogeneity in bulk rock analyses. Thegeochemical characteristics of extremely alteredcompositional domains may help decipher theorigin of volatiles or melts that may be related tothe subduction of oceanic crust.

Recycling of oceanic crust into the Earth’smantle may profoundly influence the uranium

Discussion 531


budget, and the evolution of U/Th/Pb isotopesystematics in the mantle (Hart and Staudigel,1989; Elliott et al., 1999). Uranium is readilytaken up by oceanic crust during hydrothermalalteration and recycled into the mantle whilethorium concentrations remain relativelyunchanged during ocean-crust alteration. Theuptake of uranium in the oceanic crust is restrictedto the upper alteration zones, having relativelyhigh water–rock ratios. There, oxidizing seawaterenters the crust and loses its dissolved oxygenfrom reactions with the highly reducing oceaniccrust. Most of the uranium dissolved in seawater isin the oxidized form of UO2(CO3)2

22 (i.e., U6þ),and is reduced to U4þ, which is as insoluble inhydrous solutions as thorium (Langmuir, 1978).This fixation of uranium in reducing hydrothermalenvironments is critically dependent on itsmobilization by oxidation under the present-dayatmospheric conditions. However, the earth’satmosphere has only been oxidizing for the last2.2 Gyr (Holland, 1984, 1994), and for this reason,uranium recycling is unlikely to have occurredprior to 2.2 Ga. This process may explain the“kappa-conundrum,” whereby MORB and theupper mantle appears to have much lower232Th/238U ratios (¼“kappa”) than required bymodeling of lead isotope ratios (Elliott et al., 1999and references therein). These present-day lowkappas may be caused by the recycling of uraniumrelatively recently in Earth’s history.

One of the major fluxes associated with therecycling of oceanic crust involves water andCO2. Most of the volatile inventory of alteredoceanic crust is located in the uppermost 600 m,which is also the section first exposed to thetop-down heating of the slab during subduction.For this reason, extraction of volatiles duringsubduction is particularly efficient in the upperpart of the slab and much of this inventory is likelyto be extracted (e.g., Kerrick and Connolly, 2001).However, there are several mechanisms that couldallow these elements to survive passage throughthe “subduction zone filter (see also Chapter3.17).” (i) In particular old and dense crust may besubducted relatively rapidly, which greatlyreduces the geotherms in the subducting slab,increasing volatile subduction (Staudigel andKing, 1992). (ii) Uneven topography on thesurface of subducting upper oceanic crust (dueto horst-graben structures and collapsed sea-mounts) can produce vertical throw of highlyaltered materials to levels substantially below theaverage top of the remaining slab, thus isolatingthese materials from top-down heating. (iii) Somevolatiles from dehydration may be retained asfluid inclusions in newly formed minerals thatremain stable past the subduction zone, and maymake up several percent of a rock (Touret andOlsen, 1985).

3.15.6 CONCLUSIONS

Seafloor hydrothermal alteration processes areimportant for the global geochemical cycles ofmany elements, and the record of these processesin the oceanic crust reveals much informationabout these cycles. Rather robust flux informationcan be obtained from a variety of elements thathave rather low initial abundances in basalt (H2O,CO2, K2O, rubidium, caesium, uranium) or thatare rather sensitive to alteration (87Sr/86Sr) andd18O. Fluxes of many other elements are ratherpoorly constrained because of substantial primarymagmatic variation.

This rock record yields results that are incon-sistent with fluxes inferred from fluid data fromseafloor hydrothermal springs. An in-depth ana-lysis of data methods and uncertainties in fluxesbased on rock data and fluid data suggest thatthese discrepancies are due to profound problemswith both types of flux determination.

Fluxes from fluid data are fundamentallylimited by their near-exclusive focus on reac-tions involved in large fluid fluxes, high-temperature reactions and in particular inyoung crust near the ridge axis. A few studieshave recently begun exploring low-temperaturevents in crust up to 3.5 Myr old. Hydrothermalfluxes up to this age involves only about half ofthe total convective heat flow lost in the oceans,and by far the largest volume of fluids passesthrough oceanic crust older than 3.5 Ma.Furthermore, much, if not most of the oceaniccrust is altered at low fluid/rock ratios, andthus, most of the alteration of the oceanic crustis not accounted for by in these fluxes.

Data from the rock record are limited to therather small number of drill sites that are in crustold enough and have sufficiently high recoveryrates to allow reliable estimates to be made. Noneof the drill holes available so far reaches into thereaction zone of a black smoker, which is wheremuch of the hydrothermal flux data derive from.Most importantly, the uncertainties of someelement fluxes (i.e., silicon, aluminum, mag-nesium, etc.) derived from oceanic crust studiesare substantially larger than the fluxes expectedfrom balancing other flux data. This is due to therather high concentrations of these elements in theoceanic crust, their large magmatic variation, andtheir complex behavior during alteration. It is thusunlikely that the current approach will yield fluxesfor these elements that are sufficiently constrainedto be meaningful in the context of global fluxes inthe hydrosphere.

There is no simple solution to “fix” the intrinsicproblems of fluid or rock-based estimates ofhydrothermal fluid fluxes. For this reason, bothmethods have to be used in a complementaryfashion. The first goal in such a complementary


analysis is to develop a reference model for ocean-crust hydrothermal alteration, with a clear defi-nition of reservoirs, reaction zones, types ofalteration, etc. The goal of such a referencemodel is to define the portions of oceanic crustthat can be constrained by various methods.Fluxes derived from studies of ophiolites need tobe used to constrain those of the deeper oceaniccrust, until all major fractions of the oceanic crustcan be reliably recovered by drilling. Above all,studies should evaluate uncertainties and place allflux determinations in this context.

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