Hydrothermal fluid fluxes calculated from the isotopic mass balance of thallium in the ocean crust

Earth and Planetary Science Letters 251 (2006) 120 – 133 www.elsevier.com/locate/epsl

Hydrothermal fluid fluxes calculated from the isotopic mass balance of thallium in the ocean crust Sune G. Nielsen a,b,⁎, Mark Rehkämper a,c , Damon A.H. Teagle d , David A. Butterfield e , Jeffrey C. Alt f , Alex N. Halliday a,g a

Department of Earth Sciences ETH Zurich, Sonneggstrasse 5, 8092 Zurich, Switzerland GEMOC, Department of Earth and Planetary Sciences, Macquarie University, 2109 NSW, Australia c Department of Earth Science and Engineering, Imperial College, London SW7 2AZ, United Kingdom National Oceanography Centre, School of Ocean and Earth Science, University of Southampton, SO14 3ZH, United Kingdom e University of Washington and NOAA, Seattle, WA, USA f Department of Geological Sciences, University of Michigan, Ann Arbor, MI, USA g Department of Earth Sciences, University of Oxford, Parks Road, Oxford, OX1 3PR, United Kingdom b

d

Received 9 February 2006; received in revised form 23 August 2006; accepted 1 September 2006 Available online 11 October 2006 Editor: H. Elderfield

Abstract Hydrothermal fluids expelled from the seafloor at high and low temperatures play pivotal roles in controlling seawater chemistry. However, the magnitude of the high temperature water flux of mid-ocean ridge axes remains widely disputed and the volume of low temperature vent fluids at ridge flanks is virtually unconstrained. Here, we determine both high and low temperature hydrothermal fluid fluxes using the chemical and isotopic mass balance of the element thallium (Tl) in the ocean crust. Thallium is a unique tracer of ocean floor hydrothermal exchange because of its contrasting behavior during seafloor alteration at low and high temperatures and the distinctive isotopic signatures of fresh and altered MORB and seawater. The calculated high temperature hydrothermal water flux is (0.17–2.93) × 1013 kg/yr with a best estimate of 0.72 × 1013 kg/yr. This result suggests that only about 5 to 80% of the heat available at mid-ocean ridge axes from the crystallization and cooling of the freshly formed ocean crust, is released by high temperature black smoker fluids. The residual thermal energy is most likely lost via conduction and/ or through the circulation of intermediate temperature hydrothermal fluids that do not alter the chemical budgets of Tl in the ocean crust. The Tl-based calculations indicate that the low temperature hydrothermal water flux at ridge flanks is (0.2–5.4) × 1017 kg/yr. This implies that the fluids have an average temperature anomaly of only about 0.1 to 3.6 °C relative to ambient seawater. If these low temperatures are correct then both Sr and Mg are expected to be relatively unreactive in ridge-flank hydrothermal systems and this may explain why the extent of basalt alteration that is observed for altered ocean crust appears insufficient to balance the oceanic budgets of 87Sr/86Sr and Mg. © 2006 Elsevier B.V. All rights reserved. Keywords: Ocean crust; Hydrothermal fluids; Fluid fluxes; Thallium; Basalt alteration

⁎ Corresponding author. Present address: Department of Earth Sciences, University of Oxford, Parks Road, Oxford, OX1 3PR, United Kingdom. E-mail address: [email protected] (S.G. Nielsen). 0012-821X/$ - see front matter © 2006 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2006.09.002

S.G. Nielsen et al. / Earth and Planetary Science Letters 251 (2006) 120–133

1. Introduction The high-temperature (high-T) hydrothermal circulation that occurs at mid-ocean ridge crests has a profound effect on the global geochemical budgets of numerous elements and isotope systems, including the alkalis, alkaline earths, and 87Sr/86Sr [1]. As a consequence, such hydrothermal systems have been intensely studied in the last 25 yr [2–4], although little consensus has been reached as to the magnitude of the water fluxes. Some recent investigations have furthermore suggested that the cooler hydrothermal fluids of mid-ocean ridge flanks may also generate significant chemical fluxes for some elements [5], but the geochemical significance of these low-T systems is disputed [6,7]. Simple thermal calculations of hydrothermal power outputs presently provide the most useful information on the overall magnitude of hydrothermal circulation at ridge axes and flanks [1,8]. Conversion of the power outputs to hydrothermal water fluxes and elemental fluxes is difficult, however, particularly because of the unknown partitioning of heat between (a) hot, near critical point black smoker fluids and lower temperature diffuse emissions at ridge axes and (b) warm (N 45 °C) and cool (b 20 °C) hydrothermal circulation at ridge flanks [5]. These distinctions are of great significance, due to the different chemical reactivity of hot, warm and cold vent fluids. In principle, isotope and element abundance data for hydrothermal fluids and altered ocean crust can be used to obtain independent geochemical estimates of hydrothermal water fluxes at both ridge axes and flanks. Previous work has shown, however, that only very few elements and isotope systems are able to provide useful constraints, particularly estimating the magnitude of low-T hydrothermal circulation. Here we show that the thallium (Tl) concentrations and isotope compositions of hydrothermal vent fluids, fresh mid-ocean ridge basalts (MORB) and altered ocean crust, as sampled primarily by rocks from ODP Hole 504B (which provides one of the most complete in situ section of oceanic crust [9]), can be used to constrain both axial high-T and off-axis low-T hydrothermal water fluxes. 2. Methods 2.1. Sample preparation Powdered samples of altered basalts (0.1–2 g) were dissolved in a 1:1 mixture of concentrated HF and HNO3 on a hotplate overnight. The samples were then

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dried, re-dissolved in 6 M HCl, and the resulting solutions diluted with 18 MΩ-grade water to ∼ 1 M HCl. In some cases a minor residue, which probably consisted of fluorides, persisted at this stage. A previous study has shown that such residues do not compromise the Tl contents and that they do not fractionate Tl isotope compositions [10]. For the high-T hydrothermal fluids, separate aliquots were weighed out for the determination of Tl isotope compositions and Tl concentrations (by the isotope dilution technique). The low Tl contents of the low-T fluids necessitated that the complete sample was spiked for the determination of the Tl abundance. All fluid samples were dried down after weighing, re-dissolved in aqua regia to oxidize any S present, dried again and then taken up in 1 M HCl. 2.2. Determination of Tl isotope compositions A two-stage column chromatographic technique with anion-exchange resin was used to isolate Tl from rock samples [10,11], whereas a single mini-column sufficed for the separation of Tl from the hydrothermal fluids. Total procedural Tl blanks during this study were b3 pg, which is insignificant compared to the Tl present in the samples. The Tl isotope compositions were determined at the ETH Zurich with a Nu Plasma multiple collector inductively coupled plasma-mass spectrometer (MCICPMS) using previously described techniques that utilize both external normalization to NIST SRM 981 Pb and standard-sample bracketing for mass bias correction [11,12]. Thallium isotope compositions are reported relative to the NIST SRM 997 Tl standard in parts per 10,000 such that ε205 Tl ¼ 10; 000xð205 Tl=203 Tlsample −205 Tl=203 TlSRM997 Þ ð1Þ =ð205 Tl=203 TlSRM997 Þ The uncertainty of the analyses is ± 1 ε205Tl-units (2σ) for all samples, which has been verified in previous studies [10,11]. The absence of any systematic bias was also established in earlier investigations [11]. 2.3. Determination of Tl concentrations For the hydrothermal fluids, the Tl concentrations were determined by isotope dilution (ID), whereby the samples were spiked with an enriched tracer of 203Tl [11]. The ID analyses were conducted by MC-ICPMS, using added Pt and Pb for the external normalization of mass discrimination [11]. The uncertainty of the ID analyses is about ± 1.5% (1σ) [11].

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The Tl concentrations of the rock samples were determined by monitoring the 205 Tl signal intensities of the samples during the isotopic measurements, using internal normalization to added 208Pb. These data are estimated to be accurate to about ± 25% [13]. 3. Results The results of this study confirm previous investigations [14] in finding that the Tl concentrations of endmember high-T vent fluids (∼7.6 to 58 nmol/kg) are more than two orders of magnitude higher than seawater (∼65 pmol/kg) (Table 1). The elevated Tl contents of the vent fluids result from leaching of the ocean crust at high temperatures, as evidenced by a mean Tl isotope composition for black smoker-type hydrothermal fluids (ε205Tl = −1.9 ± 0.6, 1σ) that is indistinguishable from fresh MORB glasses (ε205Tl = −1.9 ± 0.6, 1σ) (Table 2) [15] but markedly different from seawater (ε205Tl ≈ −6) (Table 1).

Large systematic variations of Tl isotope compositions and concentrations with increasing depth are apparent for the samples from ODP Hole 504B (Fig. 1, Table 2). In the volcanic zone (VZ), which forms the upper 600 m of the basement at this site, the lavas are partially altered at low temperatures (b150 °C) and have ε205 Tl-values and Tl abundances that range from about − 2 to − 16 and about 1 to 250 ng/g, respectively. It is notable that the samples with the highest Tl concentrations also yield the most fractionated Tl isotope compositions, relative to MORB. An even higher Tl concentration (∼ 500 ng/g) together with ε205 Tl ≈ − 7.5 was determined for a sample of highly altered uppermost ocean crust from DSDP Hole 417D (Table 2). The two samples from the transition zone of Hole 504B (∼ 600 to 800 msb; meters sub-basement) have intermediate Tl contents (5 to 10 ng/g) and ε205Tl ≈ − 2 to 0. In the sheeted dike complex (800 to 1800 msb), which is partially altered to greenschist facies assemblages, the

Table 1 Thallium isotope compositions and Tl, Mg, Cs, Rb and Cl concentrations of high-and low-temperature hydrothermal fluids and seawater Sample

Locationa

Temp.

ε205Tl

(°C)

Tlb

Mg

Cs

Rb

Cl

(nmol/kg)

(mmol/kg)

(nmol/kg)

(mmol/kg)

(mmol/kg)

7.60 20.0 11.7 13.1 8.07 17.8 15.2 22.5 35.0 16.1 58.1

1.55 2.53 3.15 – 7.01 4.97 1.99 1.32 1.7 3.16 2.22

47.4 97.2 37.2 – 14.0 59.2 65.7 231 – – 336

6.76 13.6 5.82 – 2.42 9.55 9.44 19.9 – – 29.9

High-temperature hydrothermal fluids 3324d10c 14°S South EPR 3325m16 14°S South EPR 3328m10 17°S South EPR 3331d1c 18.4°S South EPR 3335d13 21.5°S South EPR 3338m4 31°S South EPR 3338m10 31°S South EPR 3573m12 Bastille Vent JFR R739b16 Inferno Vent JFR R739b18 Casper Vent JFR J2-010m14 Hulk Vent JFR

380 309 366 369 404 340 340 379 292 262 347

− 1.2 – − 1.7 − 1.4 − 1.7 − 1.5 − 1.4 − 2.5 − 2.3 − 2.3 − 3.3

Off-axis low-temperature fluids J2-005bf 9 Baby Bare JFR J2-005bf 17 Baby Bare JFR J2-005pf 23 Baby Bare JFR J2-007bf 9 ODP Hole 1026B

18.3 10 18 62.2

– – – –

0.0439 0.0482 0.0439 0.222

2.29 2.93 3.92 5.16

6.45 6.67 5.81 7.60

1.09 1.17 1.14 1.13

53.2e – – – –

2.2e

− 5.8 − 5.7 − 6.1 − 6.1

0.065d 0.0503 0.0530 0.0524 0.0502

1.4e – – – –

Seawater BZ St. 7, 10 m BZ St. 7, 600 m BZ St. 7, 1200 m BZ St. 7, 2500 m

– – – –

– – – –

Cs/Rbc (×10− 3)

Tl/Clb,c (×10− 9)

141.1 486.4 431.0 662.8 354.5 524.6 523.4 379.1 539.0 311.7 509.4

7.01 7.15 6.39 – 5.77 6.20 6.64 11.65 – – 11.25

59.3 41.4 27.6 19.1 24.9 34.2 29.1 60.1 65.1 54.3 114.4

544.0 550.6 552.1 542.3

5.91 5.69 5.09 6.75

0.0807 0.0875 0.0795 0.409

546e – – – –

1.6e – – – –

0.12e – – – –

– Not determined. a EPR: East Pacific Rise; JFR: Juan de Fuca Ridge. b The Tl concentrations and Tl/Cl ratios of the axial hydrothermal fluids are endmember values calculated by extrapolation to zero Mg, except for 3331d1c for which no Mg data are available. These corrections are minor as all fluids have [Mg] ≤ 7 mmol/kg. c Elemental ratios are molar. d The average Tl concentration of seawater is from [42]. Mg, Cs, Rb and Cl data on hydrothermal fluids from Butterfield, pers. comm. e The Mg, Cs, Rb and Cl concentrations of seawater are from [55].

S.G. Nielsen et al. / Earth and Planetary Science Letters 251 (2006) 120–133 Table 2 Thallium concentrations and isotope compositions of MORB glasses and altered ocean crust from ODP Hole 504B and DSDP Hole 417D Sample

Location

MORBa SO157 54DS1

Pacific– Antarctic Rise POS 221 626 DS Kolbeinsey Ridge TT 152-21 Juan de Fuca Ridge ALV 731-4 Galapagos Ridge CH 98 DR11 MidAtlantic Ridge

ODP Hole 504B 013R4 019-023 red 013R4 019-023 dark grey 028R2 050-053 red 028R2 038-040 dark grey 042R2 047-051 dark grey 047R2 066-070 breccia 060R1 087-089 dark grey 081R1 083-086 light 082R1 090-096 light 104-2,57-61 light halo 104-2,67-72 dark grey 150R-1,138-140 dark grey 209 R1, 35-41B light patch 209 R1, 35-41A dark grey 247R1 30-33 dark grey 247R1 46-50 light halo

Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift Costa Rica Rift

DSDP Hole 417D 50-2 42-44 Bermuda dark grey Rise

Tl

Cs

(ng/g)

(ng/g)

ε205Tl

Depth subbasement (m)

7.4

–

−1.6

–

7.1

–

−2.5

–

9.7

–

−2.5

–

5.5

–

−1.8

–

4.7

−

−0.9

–

90

−14.7

91.7

59.7

10

−15.5

91.7

9.2

840

−9.4

202.5

32.5

15

−9.7

203.9

0.49

22

–

320.5

4.1

23

−1.7

365.7

3.7

19

−5.3

467.9

10.7

32

−2.0

645.9

4.5

12

−0.7

654.9

247

3.7

–

−3.4

844.0

6.4

–

−1.9

844.2

0.36 –

−2.3

1155.6

0.21 –

−1.3

1513.4

0.52 –

−1.5

1513.4

0.27 –

−2.2

1782.5

0.18 –

–

1782.7

500

–

−7.4

MORB analyses from [15].

majority of samples have low Tl abundances (b 1 ng/g) and all haveε205 Tl values of between − 1.3 and − 3.4. The pattern of Tl distribution in the ocean crust is summarized in Fig. 2 and is akin to that observed for the heavy alkali elements Rb and Cs [16,17], with high Tl concentrations for rocks of the volcanic zone and low abundances in the altered sheeted dike complex. Although Tl/Cs (and Tl/Rb) ratios are highly variable in the volcanic zone (Table 2, [18]), implying distinct deposition mechanisms for Tl and the alkalis, our data confirm that Rb, Cs, and Tl display a broadly similar geochemical behavior in hydrothermal systems [14,19,20]. All three elements are deposited in the upper ocean crust from circulating seawater at temperatures of less than ∼ 150 °C, whereas they are leached from the rocks of the sheeted dike complex by hot (N250–300 °C) hydrothermal fluids. 4. Axial high-temperature fluid flux As more than 99% of the Tl-budget of high-T vent fluids originates from the leaching of ocean crust, we can use the following mass balance equation to determine the high-T hydrothermal water fluxes of ridge-crests (FhT): FhT  ½TlhT ¼ Foc leach  ½Tloc  fTl

leach

ð2Þ

Here [Tl]hT is the average Tl concentration of the vent fluids,Foc leach is the mass flux of ocean crust that is leached by high-T fluids, [Tl]oc is the Tl content of the crust prior to leaching, and fTl leach is the fraction of Tl leached from the rocks during alteration.

217

– Not determined, red — oxidised, light — more altered, dark grey — less altered. a

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Fig. 1. Thallium concentrations (triangles) and isotope compositions (black squares with error bars) versus depth for samples from ODP Hole 504B. The dark and light shaded areas denote the average Tl concentration and isotope composition of MORB (see text). Note the change of scale for the Tl concentrations higher than 40 ng/g.

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Fig. 2. Cross-section of ocean crust that summarizes the chemical behavior of Tl during hydrothermal alteration. Shown are the measured (and in the case of low-T flank fluids, the inferred) Tl concentration and isotopic data for rocks from ODP Hole 504B, hydrothermal fluids, MORB and seawater. Arrows indicate hydrothermal circulation. msb — meters sub-basement.

This mass balance equation is, however, difficult to solve if a significant fraction of the Tl in high-T fluids derives from the VZ. The addition of Tl to the VZ rocks by alteration at low temperature generates highly variable Tl abundances and this makes it difficult to reliably establish the concentration [Tl]oc and the fraction leached,fTl leach (Table 2). It was previously concluded by Palmer and Edmond [21] that the elevated Cs/Rb ratios (up to Cs/ −3 Rb = 17 × 10 ) of high-T vent fluids from the MidAtlantic Ridge and some East Pacific Rise samples were diagnostic for the assimilation of secondary clay minerals in the volcanic zone during fluid circulation. Both elements were therefore deemed to be unsuitable for constraining global hydrothermal water fluxes. As Tl, at least to some degree, behaves similarly to Rb and Cs in hydrothermal systems it is reasonable to draw the same conclusion for Tl. However, the stark contrast in the Tl concentrations and isotope compositions of pristine MORB and low-T altered basalts allows us to screen the vent fluid data for samples that contain a significant component of Tl derived from lavas previously altered at

low temperatures. Thus, we are able to determine the ridge-crest flux of high-T hydrothermal fluids (FhT) without the influence of recycled Tl from the VZ. 4.1. Assessment of modeling parameters The calculation of a reliable Tl-based flux for high-T hydrothermal fluids requires a thorough assessment of the parameters of Eq. (2). In the following we evaluate the magnitude and uncertainty of each of these variables, which are summarized in Table 3. The uncertainties quoted for each of the parameters are based on the range of values found in the literature combined with those determined in this study. In all cases, except for the mass flux of ocean crust that was taken from [8], we have been conservative and included the entire range of values. This approach was chosen because published Tl data are scarce, such that standard deviations of data populations are unlikely to provide meaningful uncertainties for our estimates of global geochemical budgets. A conservative methodology was also adopted for establishing

S.G. Nielsen et al. / Earth and Planetary Science Letters 251 (2006) 120–133 Table 3 Parameters used for estimating the high-T vent fluid flux of mid-ocean ridge axes Parameter

Best estimate

Tl concentration of high-T vent fluids Tl concentration of MORB Tl concentration of cumulate crust Fraction of Tl leached from ocean crust Fraction of cumulates in ocean crust Mass flux of ocean crust, non-cumulate Mass flux of ocean crust, cumulate

22 ± 11 nmol/kg 3.0 ± 1.5 ng/g 0.75 ± 0.38 ng/g 0.75 ± 0.15 0.6 1.25 ± 0.17 × 1016 g/yr 0.76 ± 0.10 × 1016 g/yr

the uncertainties of the fluid flux calculations, which sum all of the quoted errors instead of using the root of the sum of squares. 4.1.1. Tl concentration of high-T vent fluids A simple method for determining the global average metal concentrations of high-T fluids is based on an evaluation of the typically systematic Metal/Cl ratios exhibited by various individual samples. Chlorine behaves conservatively in hydrothermal systems and its concentration in vent fluids varies primarily due to the brinevapor phase separation that occurs at high temperature. The Tl/Cl ratios of the present samples and of fluids analyzed in previous studies are evaluated in this way in Fig. 3. This plot shows that the majority of the vent fluids have Tl/Cl values of between 20 and 60 × 10− 9. Only two samples from the Juan de Fuca Ridge display Tl/Cl N 80 × 10− 9 and these fluids may contain some recycled Tl. This interpretation is supported by the observation that one of these fluids (J2-010m14) also displays a low ε205Tl value of −3.3 and a high Cs/Rb ratio of 11 × 10− 3 (no Tl isotope or Cs/Rb data are available for the second sample). An investigation of the available Cs/Rb data for samples plotted in Fig. 3 shows that all vent fluids from the East Pacific Rise and most Juan de Fuca Ridge samples have Cs/Rb ratios of (5.8–7.9) × 10− 3, which is slightly lower than or identical to fresh MORB (8.1 ± 0.2 × 10− 3; [22]). In contrast, a black smoker fluid from the TAG hydrothermal field (Mid-Atlantic Ridge, [14]) and two samples from the Juan de Fuca Ridge (3573m12, J2-010m14; Table 1) have Cs/Rb ≥ 11 × 10− 3, and such high values could be indicative of alkali recycling. Recycling of Tl, however, can only be confirmed for a single sample (J2-010m14, see above) because the others do not have unusually high Tl/Cl and 3573m12 has a MORB-like Tl isotopic composition (ε205Tl = −2.5). No Tl isotope data are available for TAG fluids. For samples withε205Tl N −2.5, a significant contribution of recycled Tl is very unlikely because (a) the Tl-rich rocks of the

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upper volcanic zone haveε205Tl values of between −10 and −16, and leaching of such rocks would yield fluids with lowε205Tl; (b) the rocks of the lower volcanic zone and the transition zone have less extreme Tl isotope composition but they show little or no enrichment of Tl relative to pristine MORB. It is furthermore conceivable that the higher Cs/Rb ratios of at least some vent fluids may reflect preferential leaching from the sheeted dikes rather than recycling of Cs derived from altered volcanic zone rocks. This conclusion follows from the very low Cs/ Rb ratios that have been determined for the sheeted dike complex of ODP Hole 504B [23]. In summary, we infer that the data of Fig. 3, for all samples but two, provide a mean global Tl/Cl ratio for high-T hydrothermal fluids unaffected by Tl recycling of 40 ± 20 × 10− 9. Normalization of this ratio to a chlorinity of 556 mmol/kg, which is the salinity of seawater corrected for the slight loss of water that occurs during alteration [following 8], yields an average Tl content of 22 ± 11 nmol/kg for high-T vent fluids. 4.1.2. Mass flux of ocean crust affected by high-T vent fluids The mass flux of ocean crust from which Tl is leached by high-T fluids is based on a production rate of 3.3 ± 0.2 km2/yr [24], a thickness of 6.50 ± 0.75 km [25] and a crustal density of 2.8 g/cm3. This yields a total ocean crust production rate of 6.0 ± 0.8 × 1016 g/yr [8]. However, high-T fluids only leach a fraction of the oceanic crust. Here, we utilize the hydrothermal

Fig. 3. Tl/Cl ratios of endmember hydrothermal fluids versus Cl contents. Squares represent samples from the Juan de Fuca ridge, triangles samples from the East Pacific Rise and the Mid-Atlantic Ridge. The filled symbols denote samples analyzed in the present study, the open symbols are for results from the literature [14]. Excluding two outliers with Tl/Cl N 80 × 10− 9, the data define a mean Tl/Cl ratio of (40 ± 20) × 10− 9 for endmember hydrothermal fluids.

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circulation model of Mottl [8], to constrain the depth to which the ocean crust is affected by high-T fluid circulation at ridge axes. Even though this model involves some simplistic and controversial assumptions, we adopt these parameters in the following because Mottl's estimates appear to be more reasonable than those of previous studies, and because this approach permits a direct comparison of his results with our independent geochemical estimates of hydrothermal power outputs and water fluxes. We also adopt the same constant proportions between extrusive rocks (VZ and transition zone combined), sheeted dikes and lower ocean crust of 0.7:1.2:4.6 that were applied by Mottl [8]. The calculations thus assume that the high-T hydrothermal circulation of medium and fast spreading ridges (which produce 79% of all ocean crust or 4.74 ± 0.63 × 1016 g/yr [26]) only affects 1.2 km of sheeted dikes, because deeper penetration of the fluids is hindered by the presence of a shallow magma lens. This corresponds to a mass flux of 8.8 ± 1.2 × 1015 g/yr. In contrast, the hydrothermal systems of slow-spreading ridges (21% of total crust production) are thought to extend through 5.8 km of sheeted dikes as well as lower crustal gabbros and cumulates. Cumulate rocks consistently make up about 60% of the entire ocean crust [16,27,28], which implies that 3.9 km of 4.6 km of lower oceanic crust is of cumulate origin. At slow spreading ridges, high-T fluids hence leach 11.3 ± 1.5 × 1015 g/yr of ocean crust and this includes 7.6 ± 1.0 × 1015 g/yr of cumulate rocks. Taken together, this results in a mass flux of ocean crust mass that is affected by high-T fluid leaching of 2.00 ± 0.27 × 1016 g/yr. About 62% of this flux, or 1.25 ± 0.17 × 1016 g/yr has a non-cumulate origin. 4.1.3. Tl concentration of pristine ocean crust The Tl concentrations of non-cumulate ocean crust rocks (sheeted dikes, upper gabbros) are taken to be identical to fresh N-MORB, and the compilation of Sun and McDonough [29] suggests that N-MORB have an average Tl content of 1.4 ng/g. This compilation is biased toward somewhat more depleted samples than the N-MORB data of Hofmann [30]. No Tl data are available in the latter compilation, however, but a compatible abundance can be calculated based on the Cs estimate (14.1 ng/g) and using a Tl/Cs weight ratio of 0.2 ± 0.1 [20]. This yields a mean N-MORB Tl concentration of between 1.4 and 4.2 ng/g. Based on this, our best estimate for the Tl abundance of N-MORB is 3.0 ± 1.5 ng/g. It should be noted that the fresh MORB analyses outlined in Table 2 display somewhat higher Tl concentrations than inferred for average N-MORB. The high concentrations of these samples could therefore indicate contamination

either by Fe–Mn oxyhydroxides or a slight alteration of the MORB glass. This, however, appears unlikely because the Tl isotope composition of altered MORB and Fe–Mn oxyhydroxides are about ε205 Tl = − 20 (see Section 5.1.4.) and ε205 Tl = + 13 [31], respectively, whereas the MORB samples display a very homogenous isotope composition. Whether the five analyses presented here offer a better representation of the Tl concentration of MORB than outlined above is unknown. Only a large number of additional fresh MORB analyses will clarify this issue and we have therefore chosen to use the data from the published compilations. The cumulates are thought to have highly incompatible element (including Tl) concentrations equivalent to about 25% of the non-cumulate contents [following 16]. Thus the cumulate portion of the lower ocean crust is estimated to have a Tl abundance of 0.75 ± 0.38 ng/g. 4.1.4. Fraction of Tl leached from the ocean crust The fraction of Tl that is leached from the ocean crust by high-T vent fluids ( fTl leach) can be assessed from the Tl concentrations of the sheeted dike rocks from ODP Hole 504B. The majority of these samples have low Tl contents of between about 0.2 and 0.5 ng/g (Table 2). Considering that the ocean crust of Hole 504B is depleted in incompatible trace elements by up to a factor of 3 relative to N-MORB [23], a pristine Tl concentration as low as 1 ng/g may be applicable. Based on this we infer that fTl leach ≈ 75 ± 15%. 4.2. High-T hydrothermal water flux and implications for axial power outputs Based on the values outlined above and in Table 3, Eq. (2) yields a ridge axis high-T vent fluid flux of between 0.17 and 2.93 × 1013 kg/yr, with a best estimate of 0.72 × 1013 kg/yr. This range of water fluxes is about an order of magnitude lower than previous estimates based on the marine Sr isotope mass balance (∼ 1.2 × 1014 kg/yr [32]) but it agrees very well with fluxes derived from Li isotope data for altered ocean crust and vent fluids (0.4– 2.6 × 1013 kg/yr [33]), as well as the Sr isotope profiles and sulfur budgets of rocks and anhydrite from ODP Hole 504B (0.56 ± 0.07 × 1013 and 0.47 × 1013 kg/yr, respectively [34,35]). These results underline that the significantly higher flux inferred from the 87Sr/86Sr budget of the oceans is likely to be in error. The uncertainty of the Tl-based estimate appears to be relatively large at first sight, but the approach taken in this paper has the advantage that it does not solely rely on data from ODP Hole 504B (as does [34,35]) and because the straightforward Tl isotope systematics of hydrothermal

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systems permit a more rigorous assessment of element recycling for Tl than is possible for Li. Assuming an average hydrothermal fluid temperature of 350 °C [8], the water fluxes calculated from Tl are equivalent to a ridge axis power output of between 0.08 and 1.42 TW (1012 W), with a best estimate of 0.35 TW. This result can be compared with the hydrothermal power output that is obtained from the thermal calculations of Mottl, which yield a value of 1.78 ± 0.36 TW [8]. It is evident that the results of the Tl-based calculation do not overlap with the thermal estimate. Moreover, since the uncertainty of about 14% in the ocean crust production rate propagates directly into the results for both models, their errors are correlated. For example, a low crustal production rate of 5.2 × 1016 g/yr results in a maximum Tl-based power output of 1.08 TW compared to 1.54 TW from the thermal model. Equally, the highest Tl-based value of 1.42 TW (which is obtained for a crust production of 6.8 × 1016 g/yr) should be compared to a power output of 2.02 TW from the thermal model. These examples demonstrate that there is a significant difference between the results obtained with the two methods. In the following section, we investigate whether this discrepancy can be reconciled by alternative models, which make different assumptions regarding the depth to which hydrothermal fluids circulate in the ocean crust. A minimum estimate is obtained by assuming that hydrothermal activity extends only to the base of the sheeted dike complex, whereas the maximum depth is the base of the ocean crust [8]. Fig. 4 compares the power outputs at ridge axes that are predicted by Mottl's thermal model and our Tl-based calculations, as a function of the fraction of hydrothermal fluids that penetrate the ocean crust beyond the sheeted dikes to the base of the lower crustal cumulates. The hydrothermal circulation patterns represented by Fig. 4 are very variable and the two endmember cases (no hydrothermal circulation beyond the sheeted dikes and circulation throughout the lower crust) are rather unlikely scenarios. Importantly, the data shown in Fig. 4 demonstrate that the results of the thermal and the Tl-based models do not overlap, irrespective of the extent of hydrothermal circulation. It is critical to note that the two models can only produce similar or identical results if the sheeted dikes and the lower ocean crust lose heat exclusively through high-T hydrothermal circulation. This is because the Tl-based estimates are based on the mass of seawater that is necessary to leach and transport Tl out of the sheeted dikes and lower crustal gabbros. In contrast, the thermal model accounts for the total amount of heat that is released by magma crystallization and the cooling of the new ocean

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Fig. 4. Plot of hydrothermal power outputs (a) as calculated from the thermal model of Mottl [8]) (dashed line), which is based on the heat available from the crystallization and cooling of the fresh ocean crust and (b) as derived from the Tl-based approach of the present study (full line; the dotted lines and gray shading denotes the respective error envelopes). All results are shown as a function of the average depth of the hydrothermal systems. The minimum depth of hydrothermal circulation (whereby no fluids penetrate to the base of the ocean crust) is to the base of the sheeted dike complex; the maximum extent of hydrothermal circulation is achieved when all fluids penetrate to the base of the cumulate lower ocean crust. The plot does not take into account the uncertainty of the ocean crust production rate, because this affects both models in the same manner.

crust to hydrothermal temperatures. The results of Fig. 4 therefore demonstrate that the water flux of high-T hydrothermal systems is insufficient to carry away all of the heat that is available at mid-ocean ridges crests. A detailed analysis of the data demonstrates that only about 5% to 80% (with a best estimate of about 20%) of the ridge axis power output is used to drive high-T hydrothermal circulation. Two principal mechanisms can account for the residual energy flux that is not discharged by high-T vent fluids. (a) The lower sections of the ocean crust can lose heat via intermediate-T hydrothermal circulation (as previously suggested by Teagle et al. [35]) or (b) by conduction of heat across the transition zone to the VZ. The mass balance constraints dictate that these alternative mechanisms can explain the observed heat flux discrepancy only if they are not associated with additional chemical fluxes of Tl out of the ocean crust. This is an important consideration, as they are both likely to be associated with the discharge of warm vent fluids at ridge crests. There is ample evidence for the diffuse discharge of large volumes of warm fluids at mid-ocean ridge axes.

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Such “diffuse fluids” at axial settings may either be generated by heating seawater to only moderate temperatures, such that they would constitute an energy discharge that is distinct from high-T fluids [36,37]. Alternatively, they can be produced as mixtures of highT vent fluids with seawater [38]. Only the former type of diffuse fluid, however, would be able to balance the inferred deficit in the axial power output budget. For example, an energy deficit of 1 TW can be accounted for by heating 7.9 or 1.3 × 1014 kg/yr of seawater to 10 or 60° above ambient temperatures, respectively. There is extensive evidence that some mixing between high-T hydrothermal fluids and seawater occurs to create fluids of intermediate temperature [3,38,39] but this will have no effect on our Tl-based estimates for the high-T hydrothermal water fluxes and power outputs. This statement holds true even if Tl does not behave conservatively during mixing because our calculations utilize only data for undiluted high-T vent fluids. However, such processeswill diminish the actual volume of undiluted high-T hydrothermal fluids discharged into the oceans. The Tl-based high-T fluid flux estimate should therefore be seen as a maximum value only. Additionally, many elements do not behave conservatively during vent fluid-seawater mixing [3] and therefore care is required when the high-T fluid fluxes reported here are used to obtain estimates of chemical fluxes. 5. Low-T hydrothermal water flux at mid ocean ridge flanks Hydrothermal activity is not restricted to the ridge axes and the cumulative power output through mid ocean ridge flanks is about a factor 3–4 higher than that of the axes [8,40]. As a consequence, large volumes of seawater circulate through the VZ at low temperatures and this leads to the formation of a wide variety of secondary minerals in the basalts [9]. Fig. 2 illustrates that the alteration products have elevated abundances of Tl and this Tl must ultimately be derived from circulating seawater. We can, therefore, determine the extent of low-T hydrothermal fluid circulation (FlT) by mass balance: Fvz  ð½Tlavz −½Tlpvz Þ ¼ FTl  ½Tlsw  fupt

ð3Þ

whereFvz is the mass flux of ocean crust that is affected by low-T alteration, [Tl]avz, [Tl]pvz, and [Tl]sw are the Tl concentrations of the altered volcanic zone basalts, their pristine equivalents and seawater, respectively. The fraction of Tl that is removed from seawater by basalt weathering is denoted by fupt.

5.1. Assessment of modeling parameters In the following, we discuss and justify the estimates for the different variables that are used to calculate the low-T hydrothermal water flux of ridge flanks. The results of this evaluation are summarized in Table 4. 5.1.1. Mass flux of ocean crust affected by low-T vent fluids Low-temperature hydrothermal circulation is generally thought to be restricted to the VZ [41], which is in accord with the observation that enrichments of Tl (and Rb, Cs) are only observed in the upper section of the oceanic crust. It is possible that some hydrothermal fluids of intermediate or even low temperature are also derived from below the transition zone, but the volumes of such fluids are likely to be negligible compared to the total low-T water flux. Here we use a thickness of 550 m for the VZ. When this is combined with the crustal production rate outlined in Section 4.1.2, this corresponds to a mass flux of 5.08 ± 0.59 × 1015 g/yr. 5.1.2. Tl concentration of the volcanic zone It is exceedingly difficult to determine a global average Tl content for the VZ because the rocks typically have highly variable Tl concentrations. In a previous study of Rehkämper and Nielsen [42], the VZ was estimated to have a mean Tl abundance of 120 ng/g. This value was based on Tl data for altered basalts from ODP Hole 896A [18] and recent altered MORB [19,20], in combination with Tl concentrations derived from Cs by assuming a Tl/Cs ratio of ∼ 1 [18,20] for altered basalts from DSDP Holes 417A, 417D, 418A [16] and ODP Hole 504B [23]. This value may, however, be biased towards low concentrations because the crust at ODP Holes 504B and 896A is young with an age of ∼ 6.9 Myr. In addition, the recovery rates at these ODP sites were low (b30%) and the most intensely altered lithologies may have been lost during drilling. The Tl concentrations (inferred and measured) at these two sites could

Table 4 Parameters used for estimating the low-T vent fluid flux of mid-ocean ridge flanks Parameter

Best estimate

Tl concentration of seawater Tl concentration of altered MORB Tl concentration of pristine MORB Fraction Tl removed from seawater Mass flux of upper ocean crust that is altered by low-T vent fluids

65 ± 5 pmol/kg 200 ± 150 ng/g 3.0 ± 1.5 ng/g 0.5 ± 0.2 5.08 ± 0.59 × 1015 g/yr

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therefore be unsuitable for characterizing the average Tl addition that occurs during weathering of ocean crust. More suitable samples, which record a long history of hydrothermal circulation and have higher recovery rates are from DSDP Holes 417A, 417D and 418A, and ODP Hole 801C. These drill holes are all situated in N 110 Myr old crust and they extend almost to the base of the VZ in their respective areas. The DSDP Holes 417A, 417D and 418A are located in close proximity to each other and they have been combined into a single supercomposite sample, which corresponds to the estimated weighted average of the lithologies. Two studies have determined that the supercomposite of the VZ sampled at these holes contain about 153–167 ng/g of Cs [16,43]. Using a similar approach, Kelley et al. [44] determined a Cs abundance of about 300 ng/g for the supercomposite of ODP Hole 801C. This sample, however, may be biased towards an atypically high Cs concentration because it includes significant amounts of interflow material (sediments) with high Cs contents [45,46]. As noted in Section 3, the Tl/Cs ratio of the VZ is highly variable, presumably because Tl and Cs are preferentially incorporated into different secondary minerals. This complicates the simple application of a Tl/Cs ratio of ∼ 1 for the VZ, as employed by Rehkämper and Nielsen [42]. We also adopt this approach in the present paper, however, because Tl data for VZ rocks is very

Fig. 5. Graph that summarizes the results of the Rayleigh fractionation model, which can account for the Tl concentrations and isotope compositions observed for the weathered basalts of the volcanic zone of ODP Hole 504B (see Fig. 1). Plotted are the calculatedε205Tl values of Tl deposited from circulating seawater as a function of the fraction of Tl that is taken up by the basalts ( fupt, see text for details). The bold line with dashed uncertainty envelope denotes how the isotope composition of the deposited Tl changes with fupt for fractionation factors of α = 0.9983 ± 0.0003. The thin line denotes the cumulative isotope composition of the deposited Tl.

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scarce and because no suitable alternative method for determining the average Tl concentration of the VZ is available at present. If the Cs abundances of the supercomposites are combined with Tl/Cs ∼ 1, this implies that the VZ contains about 150 to 300 ng/g of Tl. Considering the uncertainty involved in estimating the Tl/Cs ratio, we therefore infer a global average Tl abundance of 200 ± 150 ng/g for the VZ. 5.1.3. Tl concentration of seawater Rehkämper and Nielsen [42] conducted a thorough review of published Tl concentration data for seawater and arrived at a global mean value of 65 ± 5 pmol/kg. In this study, we analyzed four additional Pacific seawater samples, which display slightly lower Tl contents (Table 1). However, considering the relatively large and reliable Tl database for seawater, which includes samples with concentrations that overlap with the present results, we find no reason to revise the estimate of Rehkämper and Nielsen [42]. 5.1.4. The fraction of Tl removed from seawater during low-T alteration It is shown in the following that a simple Rayleigh fractionation model can account for the systematic changes of Tl concentrations and isotope compositions observed for the weathered volcanic zone rocks of ODP Hole 504B (Fig. 5). This model is then, in combination with Tl abundances measured for low-T fluids from the Juan de Fuca Ridge flank, exploited to estimate how much Tl is taken up by the basalts of the VZ during alteration at low temperature. The modeling assumes that seawater ([Tl]SW = 65 pmol/g with ε205Tl = − 6) is progressively depleted in Tl as it percolates downward through the basalts of the upper ocean crust. This depletion is associated with a Tl isotope fractionation of at least 9 ε205Tl-units (equivalent to a fractionation factor α = 0.9991; α = (205Tl/203Tl)alt.bas. / (205Tl/203 Tl)SW), as can be inferred by comparing the isotope compositions of the two Tl-rich samples from 90 m depth in ODP Hole 504B (ε205Tl ≈ − 15, Table 2 and Fig. 1) with seawater. The isotope fractionation that is associated with the uptake of Tl from seawater is likely to be larger, however, as these two samples probably acquired Tl from fluids that were already depleted in 203 Tl relative to seawater (and thus characterized byε205Tl N − 6), due to reactions that occurred in the upper 90 m of the volcanic zone. If we assume that the relatively depleted basalts of ODP Hole 504B [23] exhibit pristine Tl contents of ∼1 ng/g (combined withε205Tl = −2), then isotopic mass balance dictates that the Tl added to the volcanic zone

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at ∼90 m, ∼200 m, and ∼470 m depth is characterized by ε205Tl ≈ −15, −9, and −6.5, respectively. For a given fractionation factor α, we can therefore calculate fupt at these three depths. Because of the concentration distribution of Tl in the VZ of Hole 504B, fupt must be highest in the upper VZ and progressively decrease at greater depths. The most reasonable results, which reproduce this distribution of Tl in the VZ, are obtained with fractionation factors of α = 0.9983 ± 0.0003 (denoted by the bold line with dashed error envelope of Fig. 5). From Fig. 5 it is furthermore apparent that for α = 0.9983 ± 0.0003, the fractions of Tl that are removed from seawater by basalt alteration at depth intervals of 0–90 m, 90– 200 m, and 200–470 m are 0.3–0.4, ∼0.17, and ∼0.07, respectively. Based on the data obtained for the samples from the transition zone (at ∼650 m depth) it appears that only small quantities of Tl were deposited in the rocks below 470 m depth, such that the total loss of Tl from the circulating seawater is likely to be ≤0.7. Although this model is probably a gross simplification, it can readily account for both the rapid decrease in Tl concentrations and the associated increase of ε205Tl values with increasing depth in the basement of ODP Hole 504B. A minimum value of fupt ≈ 0.3 is obtained from the assumption that the majority of the low-T fluids penetrate less than 200 m into the basement, such that the volcanic zone has an averageε205 Tl ≈ − 20 (thin line of Fig. 5). The value of fupt ≈ 0.3 is also supported by three of the four analyzed low-T vent fluids from the flank of the Juan de Fuca Ridge because their Tl contents (∼ 45 pmol/kg, Table 1) are about 30% lower compared to average seawater [42]. The reason why the fourth low-temperature fluid from ODP Hole 1026 has a Tl concentration that exceeds the seawater Tl abundance by about a factor of 4 (Table 1) is unknown at present. It is possible that this increase is due to contamination of the fluid by the steel casing of the drill hole. Alternatively, it may reflect admixture of Tl-rich pore waters or high-temperature hydrothermal fluids. In addition this sample was warmer than the other three flank fluids with a temperature of 62 °C (Table 1), and this could alter the chemical reactivity of elements such as Cs and Tl. Despite their relatively low exit temperatures (Table 1), the flank fluids of this study are all relatively evolved, with low Mg concentrations of about 2 to 5 mmol/kg (Table 1). Nonetheless, fupt does not exceed 0.3 for Tl and the Cs concentrations all exceed the seawater value, which is opposite to what would be expected for fluids that reacted extensively with MORB at low temperatures. Combined, these observations indicate that the fluids experienced significant chemical interaction with basalts at temperatures higher than the

measured exit values. It is therefore unclear if the result of fupt ≈ 0.3 represents a meaningful value for such evolved fluids or low-T hydrothermal fluids in general. Nevertheless, it is considered unlikely to be merely coincidence that both the Rayleigh fractionation model and the Juan De Fuca Ridge flank fluids imply an uptake efficiency of fupt = 0.5 ± 0.2 for Tl. 5.2. Low-T vent fluid flux and implications for chemical exchange at mid ocean ridge flanks Based on the parameters of Table 4, Eq. (3) generates a low-T ridge flank hydrothermal water flux of between 0.2 and 5.4 × 10 17 kg/yr, with a best estimate of 1.5 × 1017 kg/yr. These volumes are very large and they readily exceed the riverine flux of 0.38 × 1017 kg/yr [47]. When the low-T fluid flux is combined with a ridge flank power output of 7.1 ± 2.0 TW [8], the expelled fluids are predicted to exhibit average temperature anomalies of only about 0.1 to 3.6 °C relative to ambient seawater. This result appears to be in excellent agreement with various rock alteration and flank fluid studies, which inferred that only a small fraction of the ridge flank hydrothermal power output is from warm sites, with the balance occurring at cool sites with temperatures of less than 20 °C [5,8,46]. It is, however, unknown if temperature differentials as low as 0.1 °C above ambient seawater would be able to drive fluid convection in the crust. The partitioning of temperatures can explain why the ridge-flank fluid fluxes deposit large quantities of elements that are likely to be reactive at temperatures of b10 °C in the ocean crust (e.g., Tl, Li, Rb, Cs, and U), whereas the budgets of Sr isotopes and Mg appear to be less affected on average, because these elements only become significantly more reactive as temperatures increase to N 20 °C [5,48–50]. The conclusion, that the low average exit temperatures of ridge flank fluids restricts the output fluxes of basaltic Sr (with 87Sr/86Sr ≈ 0.7025) into the ocean, is in accord with a recent isotope study of altered ocean crust, which found that flank fluids can only account for a maximum of 10% of the basaltic Sr necessary to balance the marine Sr isotope budget [7]. It therefore appears that neither axial high-T nor ridge flank low-T hydrothermal fluid fluxes are able to balance the present best estimates of riverine input of radiogenic Sr into the oceans. Interestingly, a similar situation applies to Mg. Complete removal of seawater Mg (52 mmol/kg) from axial high-T hydrothermal fluids (using the axial fluid flux calculated in Section 4.2.) will only deposit between 0.1 and 1.5 × 1012 mol/yr of Mg in the ocean crust. Based on

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a riverine Mg input of 5.5 × 1012 mol/yr [51] and maximum output fluxes to biogenic carbonates and adsorption onto clays of 1.0 × 1012 mol/yr [52], a residual Mg sink flux of at least 2.9 × 1012 mol/yr must be provided by other processes, such as low-T flank fluids. However, if the production rate of VZ rocks is 5.1 × 1015 g/yr (Table 4), it follows that the minimum amount of Mg that needs to be added to the VZ is 7.6 × 10− 4 mol/g or about 2.3% (by weight) MgO. There is no evidence that the upper ocean crust is uniformly enriched in MgO by 2.3% relative to pristine MORB as several sections of altered ocean crust were observed to display little or no detectable enrichment of Mg [43,44]. These results could, however, be biased by preferential sampling of Mg-poor lithologies, as even small losses of Mg-rich smectite or chlorite during coring would have a significant effect on the total Mg budget of the VZ. Alternatively, Mg could also be deposited in the sediments that overlay the basaltic crust through precipitation from pore fluids. The low Mg concentrations of pore waters have, however, generally been interpreted to reflect Mg-loss through reactions with the underlying basaltic crust rather than interaction with the sediments [48,53]. In summary, it appears that the chemical reactions that are associated with low-T hydrothermal activity at ridge flanks may generate non-trivial fluxes for both Sr isotopes and Mg but they also appear to be insufficient to balance the respective marine budgets. Hence, it may be necessary to re-evaluate all presently available flux data to detect possible inconsistencies. For the Sr isotope system it is conceivable, for example, that the riverine flux is unrepresentative due to the high variability of 87Sr/86Sr in catchments with different lithologies [7]. In addition, it is also possible that the oceanic budgets of both Sr isotopes and Mg are currently not at steady-state [7]. A further possibility that follows from the present work is that the diffuse axial hydrothermal fluids of intermediate temperature, which have no significant effect on the marine Tl budget, could be associated with significant fluxes of 87Sr and Mg to the lower sections of the oceanic crust. This inference is supported by seawater column studies of Ca and Mg concentrations above mid ocean ridge axes, which concluded that most of the chemical interaction at ridge crests occurs through fluids that have low or intermediate temperatures [54]. 6. Conclusions In this study we have used the isotopic mass balance of Tl in the ocean crust to determine the high- and low-T

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hydrothermal water fluxes of mid ocean ridge axes and ridge flanks, respectively. The calculations show that the high-T axial fluid flux is between 0.17 and 2.93 × 1013 kg/ yr, with a best estimate of 0.72 × 1013 kg/yr. This result is an accord with previous studies, which utilized Li isotopes as well as Sr isotope profiles and sulfur budgets of rocks and anhydrites from ODP Hole 504B [33–35], to infer that the marine Sr isotope mass balance yields highT water flux estimates (∼1.2 × 1014 kg/yr [32]) that are likely to be grossly in error. The conversion of the high-T vent fluid flux to a power output estimate reveals that only about 5 to 80%, with a best estimate of 20% of the thermal energy available at mid ocean ridge axes, is released through high-T hydrothermal activity. This indicates that the lower sections of the oceanic crust must lose heat either by conduction or through hydrothermal circulation of intermediate temperature, such that it has no significant impact on Tl budgets. Mass balance calculations for Tl in the volcanic zone imply that ridge flank hydrothermal systems have low-T water fluxes of between 0.2 and 5.4 × 1017 kg/yr, with a best estimate of 1.5 × 1017 kg/yr. Assuming a ridge flank power output of 7.1 ± 2.0 TW [8], the expelled fluids are predicted to exhibit average temperature anomalies of only about 0.1 to 3.6 °C relative to ambient seawater. If these low temperatures are correct, it appears unlikely that hydrothermal circulation at ridge flanks will be associated with sufficient Sr isotope exchange and Mg burial to balance the respective marine budgets. This conclusion is in accord with the results of previous geochemical investigations of altered ocean crust. It is possible that diffuse axial hydrothermal vent fluids of intermediate temperature may provide at least some of the missing fluxes that are required to balance the oceanic budgets of Mg and Sr isotopes. Acknowledgements We are grateful to B. Zimmermann for providing seawater samples and D. Harrison for comments on an earlier version of this manuscript. M. Meier, U. Menet, D. Niederer, B. Rütsche, C. Stirling, A. Süsli, S. Woodland, H. Williams, and the rest of the IGMR group at the ETH are thanked for keeping mass spectrometers and clean labs functioning at all times. This study was funded by grants from the ETH Zurich, the Schweizerische Nationalfonds, and the Danish Research Agency. We would also like to thank Mike Mottl and Mike Bickle for insightful reviews. This research used samples and data provided by the Ocean Drilling Program (ODP). ODP was sponsored by the U.S. National Science Foundation (NSF) and

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participating countries under the management of Joint Oceanographic Institutions (JOI) Inc.

[14]

References [15] [1] H. Elderfield, A. Schultz, Mid-ocean ridge hydrothermal fluxes and the chemical composition of the ocean, Annu. Rev. Earth Planet. Sci. 24 (1996) 191–224. [2] J.M. Edmond, C. Measures, R.E. McDuff, L.H. Chan, R. Collier, B. Grant, L.I. Gordon, J.B. Corliss, Ridge crest hydrothermal activity and the balances of the major and minor elements in the ocean — Galapagos Data, Earth Planet. Sci. Lett. 46 (1979) 1–18. [3] J.M. Edmond, A.C. Campbell, M.R. Palmer, G.P. Klinkhammer, C.R. German, H.N. Edmonds, G. Thompson, P. Rona, Time series studies of vent fluids from the TAG and MARK sites (1986, 1990) Mid Atlantic Ridge: a new solution chemistry model and a mechanism for Cu/Zn zonation in massive sulphide ore bodies, in: L.M. Parsons (Ed.), Hydrothermal Vents and Processes, Geological Society Special Publication, vol. 87, Geological society, 1995, pp. 77–86. [4] K.L. Von Damm, Controls on the chemistry and temporal variability of fluids, in: S.E. Humphris, J.E. Lupton, L.S. Mullineaux, R.A. Zierenberg (Eds.), Seafloor Hydrothermal Systems, Physical, Chemical, and Biological Interactions, Geophysical Monograph, vol. 91, AGU, Washington DC, 1995, pp. 222–247. [5] M.J. Mottl, C.G. Wheat, Hydrothermal circulation through midocean ridge flanks — Fluxes of heat and magnesium, Geochim. Cosmochim. Acta 58 (1994) 2225–2237. [6] D.A. Butterfield, B.K. Nelson, C.G. Wheat, M.J. Mottl, K.K. Roe, Evidence for basaltic Sr in midocean ridge-flank hydrothermal systems and implications for the global oceanic Sr isotope balance, Geochim. Cosmochim. Acta 65 (2001) 4141–4153. [7] A.C. Davis, M.J. Bickle, D.A.H. Teagle, Imbalance in the oceanic strontium budget, Earth Planet. Sci. Lett. 211 (2003) 173–187. [8] M.J. Mottl, Partitioning of energy and mass fluxes between midocean ridge axes and flanks at high and low temperature, in: P.E. Halbach, V. Tunnicliffe, J.R. Hein (Eds.), Energy and Mass Transfer in Marine Hydrothermal Systems, Dahlem University Press, 2003, pp. 271–286. [9] J.C. Alt, C. Laverne, D. Vanko, P. Tartarotti, D.A.H. Teagle, W. Bach, E. Zuleger, J. Erzinger, J. Honnorez, P.A. Pezard, K. Becker, M.H. Salisbury, R.H. Wilkens, Hydrothermal alteration of a section of upper oceanic crust in the eastern equatorial pacific: a synthesis of results from site 504 (DSDP legs 69, 70, and 83, and ODP legs 111, 137, 140, and 148), Proc. ODP Sci. Results 148 (1996) 417–434. [10] S.G. Nielsen, M. Rehkämper, D. Porcelli, P.S. Andersson, A.N. Halliday, P.W. Swarzenski, C. Latkoczy, D. Günther, The thallium isotope composition of the upper continental crust and rivers — an investigation of the continental sources of dissolved marine thallium, Geochim. Cosmichim. Acta 69 (2005) 2007–2019. [11] S.G. Nielsen, M. Rehkämper, J. Baker, A.N. Halliday, The precise and accurate determination of thallium isotope compositions and concentrations for water samples by MC-ICPMS, Chem. Geol. 204 (2004) 109–124. [12] M. Rehkämper, A.N. Halliday, The precise measurement of Tl isotopic compositions by MC-ICPMS: application to the analysis of geological materials and meteorites, Geochim. Cosmochim. Acta 63 (1999) 935–944. [13] M. Rehkämper, M. Frank, J.R. Hein, A. Halliday, Cenozoic marine geochemistry of thallium deduced from isotopic studies

[16]

[17]

[18]

[19]

[20] [21]

[22] [23]

[24] [25]

[26]

[27] [28]

[29]

[30]

[31]

[32]

of ferromanganese crusts and pelagic sediments, Earth Planet. Sci. Lett. 219 (2004) 77–91. S. Metz, J.H. Trefry, Chemical and mineralogical influences on concentrations of trace metals in hydrothermal fluids, Geochim. Cosmochim. Acta 64 (2000) 2267–2279. S.G. Nielsen, M. Rehkämper, M.D. Norman, A.N. Halliday, D. Harrison, Thallium isotopic evidence for ferromanganese sediments in the mantle source of Hawaiian basalts, Nature 439 (2006) 314–317. S.R. Hart, H. Staudigel, The control of alkalis and uranium in sea-water by ocean crust alteration, Earth Planet. Sci. Lett. 58 (1982) 202–212. S.R. Hart, K, Rb, Cs contents and K/Rb, K/Cs ratios of fresh and altered submarine basalts, Earth Planet. Sci. Lett. 6 (1969) 295–303. D.A.H. Teagle, J.C. Alt, W. Bach, A.N. Halliday, J. Erzinger, Alteration of upper ocean crust in a ridge-flank hydrothermal upflow zone: mineral, chemical, and isotopic constraints from hole 896A, Proc. ODP Sci. Results 148 (1996) 119–150. P.J. McGoldrick, R.R. Keays, B.B. Scott, Thallium–sensitive indicator of rock–seawater interaction and of sulfur saturation of silicate melts, Geochim. Cosmochim. Acta 43 (1979) 1303–1311. K.P. Jochum, S.P. Verma, Extreme enrichment of Sb, Tl and other trace elements in altered MORB, Chem. Geol. 130 (1996) 289–299. M.R. Palmer, J.M. Edmond, Cesium and rubidium in submarine hydrothermal fluids — evidence for recycling of alkali elements, Earth Planet. Sci. Lett. 95 (1989) 8–14. A.W. Hofmann, W.M. White, Ba, Rb and Cs in the Earth's mantle, Z. Naturforsch. 38 (1983) 256–266. W. Bach, B. Peucker-Ehrenbrink, S.R. Hart, J.S. Blusztajn, Geochemistry of hydrothermally altered oceanic crust: DSDP/ ODP Hole 504B — implications for seawater–crust exchange budgets and Sr- and Pb-isotopic evolution of the mantle, Geochem. Geophys. Geosyst. 4 (2003) art. no. 8904. B. Parsons, The rates of plate creation and consumption, Geophys. J. R. Astron. Soc. 67 (1981) 437–448. R.S. White, D. McKenzie, R.K. Onions, Oceanic crustal thickness from seismic measurements and rare-earth element inversions, J. Geophys. Res., [Solid Earth] 97 (1992) 19683–19715. E.T. Baker, Y.J. Chen, J. Phipps Morgan, The relationship between near-axis hydrothermal cooling and the spreading rate of mid-ocean ridges, Earth Planet. Sci. Lett. 142 (1996) 137–145. N.I. Christensen, M.H. Salisbury, Structure and constitution of lower oceanic crust, Rev. Geophys. 13 (1975) 57–86. C.A. Hopson, R.G. Coleman, R.T. Gregory, J.S. Pallister, E.H. Bailey, Geologic section through Samail Ophiolite and associated rocks along a Muscat-Ibra transect, Southeastern Oman Mountains, J. Geophys. Res. 86 (1981) 2527–2544. S.-S. Sun, W.F. McDonough, Chemical and isotopic systematics of oceanic basalts: implications for mantle compositions and processes, in: A.D. Saunders, M.J. Norry (Eds.), Magmatism in Ocean Basins, Blackwell Scientific, Oxford, 1989, pp. 313–345. A.W. Hofmann, Chemical differentiation of the Earth: the relationship between mantle, continental crust, and oceanic crust, Earth Planet. Sci. Lett. 90 (1988) 297–314. M. Rehkämper, M. Frank, J.R. Hein, D. Porcelli, A. Halliday, J. Ingri, V. Liebetrau, Thallium isotope variations in seawater and hydrogenetic, diagenetic, and hydrothermal ferromanganese deposits, Earth Planet. Sci. Lett. 197 (2002) 65–81. M.R. Palmer, J.M. Edmond, The strontium isotope budget of the modern ocean, Earth Planet. Sci. Lett. 92 (1989) 11–26.

S.G. Nielsen et al. / Earth and Planetary Science Letters 251 (2006) 120–133 [33] L.H. Chan, J.C. Alt, D.A.H. Teagle, Lithium and lithium isotope profiles through the upper oceanic crust: a study of seawater– basalt exchange at ODP Sites 504B and 896A, Earth Planet. Sci. Lett. 201 (2002) 187–201. [34] D.A.H. Teagle, M.J. Bickle, J.C. Alt, Recharge flux to oceanridge black smoker systems: a geochemical estimate from ODP Hole 504B, Earth Planet. Sci. Lett. 210 (2003) 81–89. [35] D.A.H. Teagle, J.C. Alt, A.N. Halliday, Tracing the chemical evolution of fluids during hydrothermal recharge: constraints from anhydrite recovered in ODP Hole 504B, Earth Planet. Sci. Lett. 155 (1998) 167–182. [36] M.J. Cooper, H. Elderfield, A. Schultz, Diffuse hydrothermal fluids from Lucky Strike hydrothermal vent field: evidence for a shallow conductively heated system, J. Geophys. Res., [Solid Earth] 105 (2000) 19369–19375. [37] G. Ravizza, J. Blusztajn, K.L. Von Damm, A.M. Bray, W. Bach, S.R. Hart, Sr isotope variations in vent fluids from 9°46′–9°54′N East Pacific Rise: evidence of a non-zero-Mg fluid component, Geochim. Cosmochim. Acta 65 (2001) 729–739. [38] J.B. Corliss, J. Dymond, L.I. Gordon, J.M. Edmond, R.P.V. Herzen, R.D. Ballard, K. Green, D. Williams, A. Bainbridge, K. Crane, T.H. Vanandel, Submarine thermal springs on the Galapagos rift, Science 203 (1979) 1073–1083. [39] T.S. Bowers, K.L. Von Damm, J.M. Edmond, Chemical evolution of Mid-Ocean Ridge hot springs, Geochim. Cosmochim. Acta 49 (1985) 2239–2252. [40] C.A. Stein, S. Stein, Constraints on hydrothermal heat-flux through the oceanic lithosphere from global heat-flow, J. Geophys. Res., [Solid Earth] 99 (1994) 3081–3095. [41] J.C. Alt, Subseafloor processes in mid-ocean ridge hydrothermal systems, in: S.E. Humphris, J.E. Lupton, L.S. Mullineaux, R.A. Zierenberg (Eds.), Seafloor Hydrothermal Systems, Physical, Chemical, and Biological Interactions, Geophysical Monograph, vol. 91, AGU, Washington DC, 1995, pp. 85–114. [42] M. Rehkämper, S.G. Nielsen, The mass balance of dissolved thallium in the oceans, Mar. Chem. 85 (2004) 125–139. [43] H. Staudigel, T. Plank, W. White, H.-U. Schmincke, Geochemical fluxes during seafloor alteration of the basaltic upper oceanic crust; DSDP sites 417 and 418, in: G.E. Bebout, D.W. Scholl,

[44]

[45]

[46] [47] [48]

[49]

[50]

[51] [52]

[53]

[54]

[55]

133

S.H. Kirby, J.P. Platt (Eds.), Subduction top to bottom, vol. 96, AGU Monograph, 1996, pp. 19–38. K.A. Kelley, T. Plank, J. Ludden, H. Staudigel, Composition of altered oceanic crust at ODP Sites 801 and 1149, Geochem. Geophys. Geosyst. 4 (2003). S. Revillon, S.R. Barr, T.S. Brewer, P.K. Harvey, J. Tarney, An alternative approach using integrated gamma-ray and geochemical data to estimate the inputs to subduction zones from ODP Leg 185, Site 801, Geochem. Geophys. Geosyst. 3 (2002). J.C. Alt, Hydrothermal fluxes at mid-ocean ridges and on ridge flanks, Comptes Rendus Geosci. 335 (2003) 853–864. A. Baumgartner, E. Reichel, The World Water Balance, Elsevier, 1975. C.G. Wheat, M.J. Mottl, Composition of pore and spring waters from Baby Bare: Global implications of geochemical fluxes from a ridge flank hydrothermal system, Geochim. Cosmochim. Acta 64 (2000) 629–642. R.M. Coggon, D.A.H. Teagle, M.J. Cooper, D.A. Vanko, Linking basement carbonate vein compositions to porewater geochemistry across the eastern flank of the Juan de Fuca Ridge, ODP Leg 168, Earth Planet. Sci. Lett. 219 (2004) 111–128. H. Elderfield, C.G. Wheat, M.J. Mottl, C. Monnin, B. Spiro, Fluid and geochemical transport through oceanic crust: a transect across the eastern flank of the Juan de Fuca Ridge, Earth Planet. Sci. Lett. 172 (1999) 151–165. E.K. Berner, R.A. Berner, The Global Water Cycle, 1987, 397 pp. J.I. Drever, The magnesium question, in: E. Goldberg (Ed.), The Sea, Marine Chemistry, vol. 5, Wiley Interscience, New York, 1974, pp. 337–358. C.R.P. Maris, M.L. Bender, P.N. Froelich, R. Barnes, N.A. Luedtke, Chemical evidence for advection of hydrothermal solutions in the sediments of the Galapagos Mounds Hydrothermal Field, Geochim. Cosmochim. Acta 48 (1984) 2331–2346. S. de Villiers, B.K. Nelson, Detection of low-temperature hydrothermal fluxes by seawater Mg and Ca anomalies, Science 285 (1999) 721–723. K.W. Bruland, Trace elements in seawater, in: J.P. Riley, R. Chester (Eds.), Chemical Oceanography, Acad. Press, London, 1983, pp. 157–221.