# Multiple parameters enable deconvolution of water-rock reaction paths in low-temperature vent fluids of the Kama ehuakanaloa (Lō ihi) seamount

Vincent Milesi, Everett Shock, Jeffrey Seewald, Elizabeth Trembath-Reichert, Sean P Sylva, Julie A Huber, Darlene S.S. Lim, Christopher R German

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## To cite this version:

Vincent Milesi, Everett Shock, Jeffrey Seewald, Elizabeth Trembath-Reichert, Sean P Sylva, et al. Multiple parameters enable deconvolution of water-rock reaction paths in low-temperature vent fluids of the Kama ehuakanaloa (Lō ihi) seamount. Geochimica et Cosmochimica Acta, 2023, 348, pp.54-67. 10.1016/j.gca.2023.03.013. hal-04755719

HAL Id: hal-04755719

https://hal.science/hal-04755719v1

Submitted on 28 Oct 2024

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1 MULTIPLE PARAMETERS ENABLE DECONVOLUTION OF WATER-ROCK

2 REACTION PATHS IN LOW-TEMPERATURE VENT FLUIDS OF THE

3 KAMA‘EHUAKANALOA (LŌ’IHI) SEAMOUNT

4 Vincent Milesi¹,²,³*, Everett Shock¹,²,⁴, Jeffrey Seewald⁵, Elizabeth Trembath-Reichert²,⁵, Sean P. Sylva⁵,

5 Julie A Huber⁵, Darlene S. S. Lim⁶ and Christopher R. German⁷

6
7 ¹Group Exploring Organic Processes In Geochemistry, Arizona State University, Tempe, Arizona, 85287,

8 USA

9 ²School of Earth and Space Exploration, Arizona State University, Tempe, Arizona, 85287, USA

10 ³Institut des Sciences de la Terre d'Orléans, Université d'Orléans, 45100, France

11 *vincentmilesi@hotmail.fr

12 ⁴School of Molecular Sciences, Arizona State University, Tempe, Arizona, 85287, USA

13 ⁵Marine Chemistry and Geochemistry, Woods Hole Oceanographic Institution, Woods Hole,

14 Massachusetts 02543, USA

15 ⁶Space Science and Astrobiology, NASA Ames Research Center, Moffett Field, California 94035, USA

16 ⁷Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts 02543,

17 USA

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21 Key words: hydrothermal vents; active intraplate volcano; low-temperature vent fluids; fluid geochemistry;

22 water-rock reactions; geochemical modeling; ocean Fe budget
23

## Highlights

24

New data for vent-fluid compositions at Kamaʻehuakanaloa including CO₂ and CH₄ concentrations

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Persistent vent-fluid composition for three decades

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Geochemical modeling of water-rock-gas reactions that produce the 20-50 °C vent fluids

27

Low-temperature basalt alteration results in high Fe/low H₂S vent fluids

28

Key role of active intraplate volcanoes on ocean Fe-cycle

29

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## Abstract

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The contribution of venting fluids at mid-ocean ridges to global ocean biogeochemical cycles is well recognized. Less is known about the role of magmatically-active intra-plate volcanoes. In this study, new compositional fluid data were acquired from 20-50 °C vent fluids at Kamaʻehuakanaloa (previously known as Lōʻihi) seamount (Hawaiʻian archipelago) and used to model the wide diversity of reaction conditions capable of producing the Fe-, Si- and CO₂-rich vent fluids observed. Our conceptual model includes a first step where seawater reacts with increasing proportions of basalt and gas as the temperature increases, and a second step where the resulting hydrothermal fluid mixes with unaltered seawater while continuing to react with basalt until the fluid mixture reaches 20 °C. A series of reaction paths were chosen to vary: the maximum temperature during Step 1 (50 to 400 °C) and the proportions of basalt and gas reacting; the degree, *F*, of low-temperature basalt alteration during Step 2, which corresponds to the extent to which the hot fluid generated during Step 1 continues to react with more basalt as it ascends to the seafloor. Our model shows that the 20-50 °C vent fluids are greatly dependent on the degree of low-temperature basalt alteration during fluid upwelling. Indeed, the compositions of Kamaʻehuakanaloa vent fluids cannot be reconciled with a general model of subsurface mechanical mixing of high-temperature end-member vent fluid and seawater alone. Instead, they require both subsurface equilibrium mixing between a ≥350 °C hydrothermal
fluid end-member and seawater and further basalt alteration that must occur as the fluid mixture rises to the seafloor. Although it involves only ~4% of the amount of basalt having reacted during Step 1, this low-temperature basalt alteration during Step 2 leads to the characteristic enrichments in Fe observed in the Kamaʻehuakanaloa vent fluids and a concomitant depletion in H₂S. We hypothesize that low-temperature basalt alteration during an extended path of fluid upwelling through the subseafloor might arise as a direct consequence of the height and steep-sloped topography of Kamaʻehuakanaloa seamount. If correct, this suggests a more general case - that input from magmatically-active intraplate volcanoes, which have been relatively overlooked throughout the history of submarine vent investigations to date, could differ significantly from global mid-ocean ridge fluxes and contribute more substantially than previously recognized to the global ocean Fe cycle.
56 1. Introduction

57 Hydrothermal circulation in the oceanic lithosphere plays a major role in global-scale ocean biogeochemical
58 budgets (Edmond et al., 1979; German and Seyfried, 2014) and in supporting microbial life at and beneath
59 the seafloor (e.g., Jannasch and Mottl, 1985; McCollom and Shock, 1997; Dick, 2019). Although
60 hydrothermal activity at mid-ocean ridges has received much attention since its discovery in the late 1970's,
61 that associated with active intraplate seamounts has been relatively overlooked (Beaulieu et al., 2015). Yet,
62 new evidence reveals that such systems may sustain major biogeochemical fluxes to the ocean that need to
63 be considered in global inventories (Bennett et al., 2011; Jenkins et al., 2020).

64 At the Kamaʻehuakanaloa seamount (previously known as Lōʻihi), the southernmost and youngest
65 submarine volcano of the Hawaiʻian hot-spot trace, it is well established that dissolved Fe released from
66 low-temperature vent fluids produces redox disequilibria that support vast microbial communities
67 dominated by iron-oxidizing bacteria (Emerson and Moyer, 2002; Emerson, 2009; Sylvan et al., 2017).
68 During the GEOTRACES era, it has become apparent that dissolved Fe from mid ocean ridge hydrothermal
69 venting can be traced long distances throughout every ocean basin (German et al., 2016). In the case of the
70 South Pacific, it has been calculated that upwelling of this Fe into high nutrient low chlorophyll surface
71 waters may account for ~30% of the drawdown of atmospheric CO2 into the Southern Ocean (Resing et al.,
72 2015). Only more recently has it been revealed, however, that dissolved Fe released from the
73 Kamaʻehuakanaloa seamount may also persist sufficiently far as it disperses through the ocean to stimulate
74 primary productivity in similar fashion in the high latitude North Pacific (Jenkins et al., 2020).

75 The Kamaʻehuakanaloa seamount is located 35 kms from the island of Hawaiʻi on the southern
76 flank of Mauna Loa volcano and rises ~ 4 km above the abyssal plain to a depth of 950 m below sea level.
77 Prior to the most recent eruptive activity in 1996, hydrothermal venting had been reported from the upper
78 portion of the south rift zone; and on the east flank of the volcano as well as at its summit (Gamo et al.,
79 1987; Sakai et al., 1987; Karl et al., 1988). After the eruption, venting was also established within a series
80 of summit pit craters (Wheat et al., 2000; Glazer and Rouxel, 2009; Rouxel et al., 2018). Except for the
one-year survey period following the 1996 eruptive activity where temperatures up to 200 °C were reported (Wheat et al., 2000), the vent fluids temperatures measured at Kama‘ehuakanaloa through a series of expeditions between 1987 and 2018 were mostly between 20 and 50 °C. At these temperatures, the vent fluids exhibit pH values from 5.5 to 6.2 and are enriched in Si, Fe, CO₂, NH₄⁺, PO₄³⁻ and Mn and depleted in H₂ and H₂S relative to mid-ocean ridge hydrothermal systems (Karl et al., 1988; Sedwick et al., 1992; Wheat et al., 2000; Glazer and Rouxel, 2009; Rouxel et al., 2018). The associated hydrothermal deposits consist of yellow-brown, well-crystallized goethite and smectite ranging in composition from Fe-montmorillonite to nontronite (Malahoff et al., 1982; De Carlo et al., 1983). This is quite distinct from basalt-hosted mid-ocean ridge hydrothermal systems where high temperature (~350 °C) black smoker fluids enriched in H₂S and variable amounts of chalcophile transition metals (e.g., Cu, Zn, Pb) result in the formation of polymetallic massive sulfide mounds (German and Von Damm, 2006).

Sedwick et al. (1992) proposed that pre-eruption Kama‘ehuakanaloa vent fluids resulted from a two-component subsurface mixing of seawater with a >350 °C hydrothermal fluid endmember produced from water-rock-gas interactions. In that interpretation, the high Fe concentrations of the vent fluids resulted from the leaching of basalt walls by the resulting CO₂-rich vent-fluid mixture during upwelling to the seafloor. But the process, as discussed by Sedwick et al. (1992), could also explain the origin of the Sr isotope signature and Si and alkali metal enrichments of the vent fluids that would otherwise serve as evidence for a hot (>350 °C) hydrothermal fluid endmember. Thus, the origin of Kama‘ehuakanaloa vent fluids and, especially, the relative contribution of low- and high-temperature water-rock interactions has never been fully resolved. Here we provide a quantitative evaluation of the water-rock-gas interactions at Kama‘ehuakanaloa to improve our understanding of the contribution that venting fluids associated with active intraplate seamounts could make to global-scale ocean biogeochemical budgets.

Specifically, we use newly collected geochemical data to calculate a series of reaction paths for water-rock-gas interactions associated with hydrothermal circulation and explore the diversity of reaction conditions capable of producing the low-temperature vent-fluid compositions observed at
Kamaʻehuakanaloa. Data were acquired during the 2018 SUBSEA expedition aboard the *E/V Nautilus* to the Kamaʻehuakanaloa seamount (Lim et al., 2019), to build upon earlier time-series and augment them with new data for dissolved CO₂ and CH₄ concentrations. Our study demonstrates how the chemical composition of low-temperature vent fluids can be used as a guide to water-rock reactions occurring at low-temperature in the shallow subsurface, and how those fluid compositions can be influenced by preceding water-rock reactions at higher temperatures. Our model results reveal a combination of parameters that can produce metal depletion or enrichment in low-temperature vent fluids which, in turn, can set export fluxes to the global ocean.

## 2. Material and Methods

### 2.1. Sampling and measurement

From August 21 to September 12, 2018, an exploratory mission to the Kamaʻehuakanaloa seamount was conducted aboard the *E/V Nautilus* as part of the SUBSEA research program. The Remotely Operated Vehicle *Hercules*, equipped with Isobaric Gas-Tight (IGT) sampling devices (Seewald et al., 2002), was used to collect duplicate fluid samples at 6 vents for a total of 12 samples. The pH (25 °C) value, the alkalinity and the CH₄ and H₂ concentrations, which are time-sensitive analyses, were measured onboard daily. A Gas Chromatograph (GC) equipped with a Thermal Conductivity Detector (TCD) and a Flame Ionization Detector (FID) was used to measure H₂ and CH₄, respectively, with detection limits of 3 µM and 0.05 µM. H₂S concentrations were below the detection limits of the analytical method of Cline (1969) available onboard ship, i.e., <1 µM. At the Woods Hole Oceanographic Institution laboratory, major cations and anions were analyzed with inductively coupled plasma mass spectrometry (ICP-MS) and ion chromatography. Concentrations reported for major cations are from ICP-MS except for Ca for which interference with ⁴⁰Ar⁺ prevents accurate measurements. GC-TCD was used to measure total dissolved inorganic carbon (DIC) as CO₂ after acidification of samples with a 25 wt.% phosphoric acid solution.

### 2.2. Modeling
Geochemical modeling was performed with the EQ3/6 geochemical software (Wolery, 2010) using thermodynamic data for minerals and dissolved inorganic aqueous species from Helgeson et al. (1978), Shock and Helgeson (1988), Shock et al. (1989, 1992, 1997) and Sverjensky et al. (1997). The Gibbs free energies for clay minerals were estimated using the method of Wolery and Jove-Colon (2004). The other thermodynamic parameters (S, V, Cp) are from Helgeson et al. (1978). All modeling was performed in the system Na₂O-Cl-MgO-CaO-K₂O-SiO₂-SO₂-CO₂-Al₂O₃-FeO-CuO-ZnO-TiO₂. MnO was excluded from the modeling because the limited number of Mn-bearing minerals in our database prevents a realistic description of Mn behavior. The model used for the calculation of activity coefficients is based on the B-dot equation for electrically charged species (Helgeson, 1969). For neutral solute species, the activity coefficient is taken equal to that of aqueous CO₂ in pure sodium chloride solutions of the same ionic strength (Garrels and Thompson, 1962; Helgeson, 1969). Using this approach, the activity coefficient of neutral solute species increases with increasing ionic strength, which is consistent with the “salting out” effect.

Although the Kamaʻehuakanaloa summit is 950 m below sea level, the subseafloor heat source could be much deeper as the volcano rises from >5000 m depth at its base. We considered a single constant pressure of 500 bars for all our calculations, which corresponds to a depth of ~5000 m below sea level. This assumption is reasonable because the pressure has relatively little effect on thermodynamic equilibria involving condensed phases and dissolved aqueous species compared to the effects of temperature. Our model assumptions hit their limits if phase separation were to occur due to a combination of high temperature (>400 °C) and low pressure (<300 bars) (Bischoff and Rosenbauer, 1985), as might happen during eruptive events (Wheat et al., 2000).

Our conceptual model of water-rock reactions during hydrothermal circulation involves two reaction steps as illustrated in Figure 1. In Step 1, we quantify the results as seawater reacts in open systems with Kamaʻehuakanaloa basalt while the temperature rises progressively and volcanic gases are added (CO₂, SO₂ and H₂O). In Step 2, the resulting high-temperature fluid is mixed with seawater in a second open system, while continuing to react with the basaltic interior of the volcano as it rises toward the summit,
until the fluid mixture reaches the temperature of 20 °C. This scenario was chosen to account for the water-rock-gas reactions proposed by Sedwick et al. (1992) that produce the Kama‘ehuakanaloa vent fluids. In order to investigate a large range of possible reaction conditions varying in terms of temperature, rock/water ratios and gas inputs, we used the programming language Python to automate the multiple runs and data processing of model results following the approach of Ely (2020) and Milesi et al. (2021).

During Step 1, the temperature increases from 2 °C to a temperature maximum (Tmax) ranging from 50 to 400 °C, at intervals of 50 °C (Figure 1a). The composition of seawater is taken from Bruland and Lohan (2006). Basalt composition at Kama‘ehuakanaloa includes tholeiitic, transitional and alkalic basalt as well as basanite (Moore et al., 1982; Dixon and Clague, 2001). On the upper portion of the south rift zone where the vent fields visited during the 2018 SUBSEA expedition are located, the basalt is moderately alkalic in composition, weakly phyric, containing 1-2 vol.% of olivine, plagioclase and augite, and becomes more tholeiitic going south (Karl et al., 1988). In our model, we used the basalt composition reported by Karl et al. (1988) at the Kama‘ehuakanaloa summit which falls within the compositional range described by Dixon and Clague (2001). The rock/water ratio (R/W) of Step 1, i.e., the mass of basalt in kilograms reacting with 1 kg of water, increases as the temperature maximum increases (Figure 1b), following the equation:

$$R/W_{Step1} = \frac{T_{max}}{350} \qquad (1)$$

There is essentially an infinite number of temperature-rock/water ratio paths, only a few of which are considered through Equation (1). After several iterations, we settled on this particular path as it is consistent with previous practices and observational data (Bowers and Taylor, 1985). Especially, Equation (1) leads to a rock/water ratio of 1 at 350 °C, consistent with literature estimates (Von Damm et al., 1985).

The Lo’ihi vent fluids are known to be impacted by magmatic degassing (Karl et al., 1988; Sedwick et al., 1992; Wheat et al., 2000); however, the timing of CO₂ and SO₂ inputs during hydrothermal circulation is unconstrained. Inputs of magmatic gas were considered to be a linear function of the primary
concentrations of volatiles in basalt and the rock/water ratio. Because CO₂, SO₂ and H₂O are volatile species that tend to degas during eruption, their primary concentrations were reconstructed using ratios of volatile to non-volatile elements that may behave similarly during melting and crystallization (Michael, 1988; Saal et al., 2002; Aubaud, 2004; Cartigny et al., 2008; Rosenthal et al., 2015). The CO₂, SO₂, and H₂O concentrations were reconstructed using Ba, Ce, and Dy concentrations, respectively (Saal et al., 2002; Rosenthal et al., 2015). Values of 133, 168 and 225 were adopted for the CO₂/Ba, H₂O/Ce and S/Dy ratios, respectively. The Ba, Ce and Dy concentrations for Kama‘ehuakanaloa basalt are from Dixon and Clague (2001) (Table S 1).

During Step 2 of the model, seawater at 2 °C is added to the hydrothermal fluid endmember until the fluid mixture reaches 20 °C (Figure 1a). Temperature is controlled only by mixing of the fluids, so that the addition of each increment of seawater lowers the temperature of the fluid mixture. As the temperature is lowered, proportionally greater amounts of seawater are required to lower the temperature further. Because the reaction progress of Step 2 is proportional to the amount of added seawater, most of the water-rock interaction occurs at low temperature, <50 °C. The rock/water ratio of Step 2 is defined by:

$$R/W_{Step2} = F * R/W_{Step1} \qquad (2)$$

where F (referred hereafter as the ‘degree of low-temperature basalt alteration’) is the fraction of basalt altered relative to Step 1 and was varied from 0 to 20% (Figure 1b). The small proportion of basalt reacting during Step 2 (F) compared to that reacting during the initial high-temperature reactions at depth (Step 1) is meant to account for rapid flow rates and temperature decreases from the temperature maximum reached during Step 1 as the fluids continue to upwell and mix. The proportion of gas reacting during Step 2 is calculated similarly as in Step 1.

Step 2 was modeled considering either equilibrium chemistry (referred to hereafter as ‘equilibrium mixing’) or non-equilibrium chemistry for which only aqueous speciation reactions proceed but redox reactions and mineral precipitation are suppressed (referred hereafter as ‘mechanical mixing’). The latter
case aims to account for possible disequilibrium between rapidly ascending high-temperature fluids, the
rock, and the added seawater. During mechanical mixing, no low-temperature basalt alteration is considered
as it would otherwise lead to unrealistically high solute concentrations. In our equilibrium calculation,
methane is suppressed because the reduction of seawater-derived CO₂ during hydrothermal circulation
within the oceanic crust is known to be kinetically inhibited (McDermott et al., 2015). Quartz and dolomite
are also suppressed because kinetic limitations are expected to prevent their formation at low temperature,
especially during Step 2 (Rimstigt and Barnes, 1980; Von Damm et al., 1991; Arvidson, 1999). After
several iterations and for consistency with observational data, nontronite was also removed from our
mineral database, leaving saponite, montmorillonite and beidellite among the minerals of the smectite-
group that can form.

For all fluid samples, *in situ* pH values were calculated with EQ3 by speciating the fluid composition
at 25 °C and 1 bar with charge balance on chloride, then speciating the resulting fluid composition at the
temperature of sampling (Reed and Spycher, 1984). Published data of vent-fluid composition are taken
from prior studies by Karl et al. (1988), Sedwick et al. (1992, 1994), Hilton et al. (1998), Glazer and Rouxel
(2009) and Rouxel et al. (2018). Sedwick et al. (1992) reported ranges of concentrations: only the maximum
and minimum values from that work are considered in this study. Sedwick et al. (1994) and Hilton et al.
(1998) reported only concentrations for dissolved silica and dissolved inorganic carbon, so no *in situ* pH
was calculated from these data.

## 3. Results

### 3.1. Measured vent-fluid compositions

The location and chemical composition of the Kama‘ehuakanaloa vent fluids sampled during the 2018 SUBSEA expedition are shown in Table 1. The composition of the vent fluid sampled at Marker 34 is compared to seawater values and to the end-member compositions of vent fluids from hydrothermal systems in a variety of geologic environments on the seafloor in Table 2. Basalt-hosted hydrothermal systems on mid-ocean ridges are represented by the 21°N East Pacific Rise (EPR) vent field, characterized
by 350°C fluids exiting black smoker chimneys (Von Damm et al., 1985). Hydrothermal systems in back-
arc basins associated with felsic volcanic rocks are represented by the Mariner vent field located along the
Eastern Lau Spreading Center. There, input of acidic volatiles (CO2, SO2, HCl and HF) in the circulating
hydrothermal fluids produces >350 °C vent fluids with a low pH value (2.7) and high metal concentrations
(Mottl et al., 2011). Ridge flank systems are represented by Baby Bare, an isolated seamount on the eastern
flank of the Juan de Fuca Ridge where low-temperature (<100 °C) hydrothermal reactions in the basement
produce 25 °C hydrothermal springs (Wheat and Mottl, 2000).

The temperature range of 15° to 41°C measured during the 2018 SUBSEA expedition is comparable
to that of Baby Bare springs (25 °C) and much lower than the temperatures that characterize vent fluids at
21°N EPR and Mariner. The pH (25 °C) values ranging from 5.4 to 6.2 at Kama‘ehuakanaloa are lower than
at Baby Bare but higher than the values of 3.4 and 2.7 in 21°N EPR and Mariner vent fluids, respectively.
One of the most striking features of the Kama‘ehuakanaloa vent fluids is the high values of alkalinity up to
14.3 meq/kg whereas alkalinity values in the 21°N EPR, Mariner, and Baby Bare vent fluids are either
negative or close to 0. In comparison to the other vent fluids, the Cl, sulfate, Na, K and Mg concentrations
at Kama‘ehuakanaloa are relatively close to seawater values. In the Spillway and Lohiau areas, the
Kama‘ehuakanaloa vent fluids are slightly depleted in Mg and sulfate but to a much lower extent that what
is observed in the 21°N EPR, Mariner, and Baby Bare vent fluids. With concentrations up to 16.3 mmol/kg,
Ca can be enriched at Kama‘ehuakanaloa relative to seawater, but the enrichment is less than at Mariner
and Baby Bare. The Si concentration up to 2.8 mmol/kg at Kama‘ehuakanaloa is 5 to 10 times higher than
seawater and Baby Bare values but it is lower than that of the >350 °C vent fluids. Similarly, Fe and Mn
are enriched compared to seawater and Baby Bare springs, with concentrations up to 0.98 and 0.07
mmol/kg, respectively. These values are 1 to 2 orders of magnitude lower than at 21°N EPR and Mariner.
CH4 at Kama‘ehuakanaloa is enriched compared to seawater but to a lesser extent than in black smoker
systems, whereas H2 and H2S are not detected. Finally, the CO2 concentration ranging from 14 to 126
mmol/kg at Kamaʻehuakanaloa is one order of magnitude higher than in seawater and comparable to values
of the Marnier vent field (Mottl et al., 2011).

New data for Kamaʻehuakanaloa vent-fluid compositions from the 2018 SUBSEA expedition are
compared to previously published data for vent fluids sampled at the summit of the volcano (Karl et al.,
1988; Sedwick et al., 1992, 1994; Hilton et al., 1998; Glazer and Rouxel, 2009; Rouxel et al., 2018) (Figure
2). This comparison indicates that overall, the composition and temperature of the vent fluids were
relatively constant through time from 1988 to 2018 except for the short-lived event in 1996. Small variations
exist for pH, Si, DIC and Fe concentrations but no trend with time is identified. The tectonic-volcanic event
of July-August 1996 created a pit crater (Pele's Pit) and several sites of vigorous hydrothermal venting,
with temperature up to 200 °C measured in September 1997 (Davis and Clague, 1998; Wheat et al., 2000);
however, subsequent exploratory dives over three cruises between 2006 and 2008 (Glazer and Rouxel,
2009; Rouxel et al., 2018) and the 2018 SUBSEA exploration reported temperatures (15-41 °C, Table 1)
and fluid compositions similar to pre-eruptive values (Karl et al., 1988; Sedwick et al., 1992, 1994).

As a hydrothermal fluid end-member mixes with seawater beneath the seafloor, the chemical
signatures inherited from the water-rock-gas reactions can be progressively lost. The degree of loss depends
on the involvement of chemical species in the water-rock-gas reactions and their concentrations in seawater.
Therefore, we prioritized the use of Si, pH, DIC, Ca and Fe to constrain the geochemical models because
these are the species that differ most from seawater values. Although CH₄ is also enriched relative to
seawater, unambiguous information is less likely to be obtained from modelling this species distribution
because its concentration can be the result of numerous processes including biogenic formation (e.g.,
Bradley and Summons, 2010), pyrolysis of organic matter (Welhan, 1988), kinetic limitations inhibiting
abiotic formation and oxidation (McDermott et al., 2015, 2020) and leaching of fluid inclusions and gas
vesicles (Klein et al., 2019; Grozeva et al., 2020).
<table><caption>Table 1. Locations and chemical compositions of fluid samples collected during the 2018 SUBSEA
expedition</caption><thead><tr><th>Sample (NA100-)</th><th>002</th><th>001</th><th>039</th><th>040</th><th>019</th><th>054</th><th>053</th><th>061</th><th>060</th><th>080</th><th>079</th><th>059</th></tr><tr><th>Marker<sup>a</sup></th><td>34</td><td>34</td><td>38</td><td>38</td><td>31<sup>f</sup></td><td>2</td><td>2</td><td>57</td><td>57</td><td>57<sup>g</sup></td><td>57<sup>g</sup></td><td>Bottom seawater</td></tr><tr><th>Site<sup>a</sup><br>(lat; long; depth (m))</th><td colspan="5">Spillway Area<br>(18.906; -155.257; 1270±3)</td><td colspan="3">Lohiau area<br>(18.909; -155.257; 1179±2)</td><td colspan="4">Pohaku<br>(18.901; -155.258; 1184±2)</td></tr></thead><tbody><tr><th>Temperature max. (°C)<sup>a</sup></th><td>41</td><td>41</td><td>39</td><td>37</td><td>31</td><td>22</td><td>21</td><td>15</td><td>18</td><td>21</td><td>21</td><td>-</td></tr><tr><th>pH (25 °C)</th><td>5.8</td><td>6.0</td><td>6.0</td><td>6.1</td><td>6.0</td><td>6.2</td><td>6.2</td><td>5.4</td><td>5.7</td><td>5.7</td><td>5.7</td><td>7.4</td></tr><tr><th>Alkalinity<sup>b</sup><br>(meq/kg)</th><td>14.3</td><td>14.3</td><td>13.5</td><td>9.2</td><td>8.4</td><td>8.8</td><td>8.9</td><td>6.5</td><td>11.6</td><td>13.6</td><td>12.9</td><td>2.3</td></tr><tr><th>Na (mmol/kg)</th><td>459</td><td>463</td><td>464</td><td>460</td><td>462</td><td>490</td><td>464</td><td>465</td><td>463</td><td>466</td><td>464</td><td>-</td></tr><tr><th>Cl<sup>b</sup> (mmol/kg)</th><td>-</td><td>545</td><td>545</td><td>542</td><td>544</td><td>541</td><td>543</td><td>540</td><td>538</td><td>537</td><td>538</td><td>539</td></tr><tr><th>Mg (mmol/kg)</th><td>51.5</td><td>51.5</td><td>52.0</td><td>52.2</td><td>50.0</td><td>52.5</td><td>52.8</td><td>54.0</td><td>55.0</td><td>54.7</td><td>54.5</td><td>-</td></tr><tr><th>Ca<sup>b</sup> (mmol/kg)</th><td>15.1</td><td>15.8</td><td>15.3</td><td>14.3</td><td>14.9</td><td>16.0</td><td>16.3</td><td>10.9</td><td>11.3</td><td>11.4</td><td>10.6</td><td>10.3</td></tr><tr><th>K (mmol/kg)</th><td>12.2</td><td>-</td><td>12.1</td><td>11.3</td><td>12.4</td><td>16.6</td><td>11.4</td><td>10.7</td><td>11.1</td><td>11.7</td><td>11.7</td><td>-</td></tr><tr><th>Sulfate (mmol/kg)<sup>c</sup></th><td>27.0</td><td>26.7</td><td>26.9</td><td>27.3</td><td>25.8</td><td>27.0</td><td>27.0</td><td>27.9</td><td>27.3</td><td>26.6</td><td>26.8</td><td>28.2</td></tr><tr><th>Si (mmol/kg)</th><td>2.8</td><td>2.7</td><td>2.7</td><td>1.5</td><td>1.9</td><td>1.7</td><td>1.5</td><td>0.6</td><td>1.4</td><td>1.8</td><td>1.9</td><td>-</td></tr><tr><th>Fe (µmol/kg)</th><td>413</td><td>517</td><td>579</td><td>289</td><td>384</td><td>369</td><td>285</td><td>308</td><td>758</td><td>980</td><td>978</td><td>0</td></tr><tr><th>Mn (µmol/kg)</th><td>40</td><td>40</td><td>38</td><td>24</td><td>25</td><td>73</td><td>45</td><td>10</td><td>18</td><td>21</td><td>22</td><td>-</td></tr><tr><th>Al (µmol/kg)</th><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>-</td></tr><tr><th>Sr (µmol/kg)</th><td>104</td><td>103</td><td>103</td><td>95</td><td>98</td><td>-</td><td>102</td><td>87</td><td>88</td><td>87</td><td>87</td><td>-</td></tr><tr><th>Ba (µmol/kg)</th><td>0.7</td><td>0.7</td><td>0.9</td><td>0.7</td><td>0.9</td><td>-</td><td>0.7</td><td>0.7</td><td>1.0</td><td>0.6</td><td>1.0</td><td>-</td></tr><tr><th>B (µmol/kg)</th><td>398</td><td>402</td><td>424</td><td>410</td><td>398</td><td>553</td><td>388</td><td>424</td><td>452</td><td>491</td><td>484</td><td>-</td></tr><tr><th>Li (µmol/kg)</th><td>53</td><td>53</td><td>52</td><td>42</td><td>44</td><td>76</td><td>40</td><td>32</td><td>38</td><td>45</td><td>45</td><td>-</td></tr><tr><th>Rb (µmol/kg)</th><td>5.0</td><td>4.9</td><td>4.7</td><td>3.3</td><td>4.6</td><td>5.8</td><td>3.5</td><td>2.0</td><td>2.8</td><td>3.4</td><td>3.4</td><td>-</td></tr><tr><th>Cs (µmol/kg)</th><td>0.24</td><td>0.24</td><td>0.26</td><td>0.23</td><td>0.26</td><td>0.48</td><td>0.24</td><td>0.24</td><td>0.24</td><td>0.24</td><td>0.25</td><td>-</td></tr><tr><th>Cu (µmol/kg)</th><td>0.9</td><td>1.1</td><td>1.4</td><td>5.3</td><td>1.6</td><td>6.4</td><td>0.5</td><td>1.6</td><td>1.3</td><td>0.9</td><td>1.8</td><td>-</td></tr><tr><th>Zn (µmol/kg)</th><td>27.6</td><td>27.9</td><td>7.7</td><td>6.0</td><td>32.5</td><td>134.4</td><td>12.4</td><td>22.9</td><td>4.3</td><td>4.0</td><td>14.2</td><td>-</td></tr><tr><th>H<sub>2</sub> (µmol/kg)<sup>d</sup></th><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>-</td></tr><tr><th>CO<sub>2</sub> (mmol/kg)<sup>d</sup></th><td>31.2</td><td>31.2</td><td>30.0</td><td>18.3</td><td>17.4</td><td>14.2</td><td>14.0</td><td>43.8</td><td>94.1</td><td>124.3</td><td>125.8</td><td>-</td></tr><tr><th>CH<sub>4</sub> (µmol/kg)<sup>e</sup></th><td>1.5</td><td>1.3</td><td>1.2</td><td>0.8</td><td>1.4</td><td>0.6</td><td>0.6</td><td>1.6</td><td>2.8</td><td>4.3</td><td>3.8</td><td>-</td></tr><tr><th>H<sub>2</sub>S (µmol/kg)<sup>f</sup></th><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>-</td></tr></tbody></table>

a Marker and site from Rouxel et al. (2018)

b Temperature maximum recorded during the IGT sampling

c Analysis with ICP-MS, unless otherwise noted: bGran titration, cion chromatography, dGC-TCD, eGC-FID, fspectrophotometry Cline (1969)

g Duplicate sample leaked

h n.d. = not detected

i Estimated uncertainties (2σ) are as follows: pH (25°C) values ±0.1 unit; Na, Mg, Cl, Ca, K, sulfate concentrations ±3%; SiO₂, Fe, Mn, Sr, Ba, Li, Rb, Cs, Cu, Zn concentrations ±5%; CH₄ and CO₂ concentrations ±10% and ±5%, respectively.

<table><caption>Table 2. Vent-fluid compositions from Kama‘ehuakanaloa compared with submarine vent fields from various other geologic settings</caption><thead><tr><th>Vent field</th><th>Seawater</th><th>Kama‘ehuakanaloa <sup>a</sup></th><th>21°N, EPR <sup>b</sup></th><th>Mariner</th><th>Baby Bare<br>(Spring water)</th></tr><tr><th>Geologic setting</th><th>-</th><th>Intraplate volcano</th><th>Mid ocean ridge</th><th>Back arc</th><th>Ridge flank</th></tr></thead><tbody><tr><th>Max. T (°C)</th><td>2</td><td>41</td><td>350</td><td>363</td><td>25</td></tr><tr><th>pH (25 °C)</th><td>7.8</td><td>5.9</td><td>3.4</td><td>2.7</td><td>8.3</td></tr><tr><th>Alk. (meq/kg)</th><td>2.3</td><td>14.3</td><td>-0.4</td><td>-2.25</td><td>0.43</td></tr><tr><th>Cl (mmol/kg)</th><td>546</td><td>546</td><td>489</td><td>605</td><td>554</td></tr><tr><th>SO₄ (mmol/kg)</th><td>28</td><td>26.9</td><td>0.5</td><td>-</td><td>17.8</td></tr><tr><th>Na (mmol/kg)</th><td>468</td><td>460.8</td><td>432</td><td>439</td><td>473</td></tr><tr><th>K (mmol/kg)</th><td>10.2</td><td>12.4</td><td>23</td><td>36.9</td><td>6.9</td></tr><tr><th>Ca (mmol/kg)</th><td>10.3</td><td>15.5</td><td>16</td><td>43.8</td><td>55.2</td></tr><tr><th>Mg (mmol/kg)</th><td>53</td><td>51.5</td><td>0</td><td>1.9</td><td>0.98</td></tr><tr><th>Si (mmol/kg)</th><td>0.1</td><td>2.8</td><td>18</td><td>15.1</td><td>0.4</td></tr><tr><th>Fe (µmol/kg)</th><td>&lt;0.001</td><td>465</td><td>1,664</td><td>12991</td><td>&lt;0.1</td></tr><tr><th>Mn (µmol/kg)</th><td>&lt;0.001</td><td>40</td><td>960</td><td>5723</td><td>7.8</td></tr><tr><th>CH₄ (µmol/kg)</th><td>-</td><td>1.4</td><td>70</td><td>-</td><td>-</td></tr><tr><th>H₂ (mmol/kg)</th><td>-</td><td>b.d.l</td><td>1.7</td><td>-</td><td>-</td></tr><tr><th>CO₂ (mmol/kg)</th><td>-</td><td>31.2</td><td>5.7</td><td>39.8</td><td>0.4</td></tr></tbody></table>

277

278

279

280

281

282

283

284

285

286
<table><tr><td>H₂S (mmol/kg)</td><td>-</td><td>b.d.l</td><td>7.3</td><td>9.3</td><td>-</td></tr></table>

a Average values of the 2018 SUBSEA fluids sampled at Marker 34.

b Measured composition of vent fluids are corrected for entrainment of seawater during sampling based on the quantitative removal of Mg from the fluid during water-rock reactions.

Data are from Bruland et al. (2006) for seawater, Humphris and Klein (2017) for 21°N EPR and Mariner vent fields (original data from Von Damm et al. (1985) and Mottl et al. (2011), respectively), Wheat and Mottl (2000) for Baby Bare.

Abbreviation: Max. T = maximum temperature; Alk. = Alkalinity; - = not determined; b.d.l = below detection limit.

## 3.2. Modeled fluid compositions

### 3.2.1. pH, Ca and DIC

Our calculations predict a wide diversity of fluid compositions over the range of Step 1 and Step 2 reaction conditions as shown in Figure 3. During Step 1, the increase of temperature and rock/water ratio results in a decrease of pH (Figure 3a). After a sharp increase, the Ca concentration decreases as well because of the precipitation of smectite, epidote and tremolite which fixes Ca (Figure 3b, Figure S 1 to Figure S 3 in supplementary information). The DIC concentration increases in parallel due to the inputs of CO₂ (Figure 3c), linearly correlated to the rock/water ratio (see section 2.2). As a result, the hydrothermal fluid is at equilibrium with calcite all along Step 1, except for temperatures above 350 °C (Figure 3d).

At the beginning of the mixing between the hydrothermal fluid and seawater during Step 2, the pH value decreases drastically (Figure 3a). H⁺ is produced by the precipitation of minerals in response to the supply of cations by seawater. As the mixing progresses and the temperature of the fluid mixture decreases, the pH values increase. Overall, the pH values of the 20 °C fluids at the end of Step 2 decrease with the increase of Step 1 temperature maximum. In addition, there is an increase of pH with the degree of low-temperature basalt alteration during Step 2 as H⁺ is consumed by basalt dissolution. This is especially pronounced in 20 °C fluids produced by reaction paths with temperature maxima >300 °C. Consequently, the diversity of pH values in the 20 °C fluids decreases with the increase of low-temperature basalt alteration. For F values >7%, all reaction paths converge toward the same pH values at the end of Step 2, regardless of the temperature maximum reached during Step 1.

As mixing with seawater progresses during Step 2, the Ca concentrations in fluids produced by reaction paths with Step 1 temperature maxima >250 °C increase whereas the DIC concentrations decrease
(Figure 3b and c). This trend is reversed for reaction paths with lower temperature maxima. Overall, the Ca and DIC concentrations in the 20 °C fluids decrease and increase, respectively, as the temperature maximum in Step 1 increases. In addition, these concentrations increase with the degree of low-temperature basalt alteration because the dissolution of basalt supplies both Ca and CO₂. Although less apparent than in the case of pH, there is also a loss of diversity for the Ca and DIC concentrations of the 20 °C fluids as low-temperature basalt alteration increases. For F values >7%, all reaction paths tend to produce similar Ca and DIC concentrations, independently of Step 1 reaction conditions.

The combined behavior of pH, Ca and DIC concentrations leads the fluid mixture to depart from equilibrium with calcite at the beginning of Step 2 before returning to saturation with calcite at high degrees of mixing at the end of Step 2 (Figure 3d). Equilibrium with calcite is especially common in 20 °C fluids inherited either from reaction paths with low Step 1 temperature maxima or high degrees of low-temperature basalt alteration. By buffering the pH value of 20 °C fluids produced by reaction paths with F values >7%, equilibrium with calcite causes a loss of chemical diversity. Therefore, the dependence of the 20 °C fluid compositions on Step 1 and Step 2 reaction conditions relies on the capacity of the system to counterbalance the pH increase that leads to calcite saturation at the end of Step 2. Two competing effects influence the evolution of pH: (1) high CO₂ supply and high-temperature mineral precipitation during Step 1 tend to lower the pH values of the resulting 20 °C fluids, whereas (2) H⁺ consumption by low-temperature basalt alteration tends to increase the pH values.

### 3.2.2. Sulfate, Mg, K, Si, Na and Cl

During Step 1, the increase of temperature and rock/water ratio result in the decrease of sulfate and Mg concentrations as sulfur is fixed into pyrite and Mg into saponite (Figure 4a and b, Figure S4 and Figure S1 in supplementary information). After a sharp decrease, the K concentrations increase, controlled by the precipitation of muscovite at high temperature (Figure 4c, Figure S5 in supplementary information). The Si concentrations increase reaching values close to equilibrium with quartz (Figure 4d, Figure S6 in supplementary information).
During Step 2, mixing with seawater supplies sulfate and Mg to the fluid mixture (Figure 4a and b). The sulfate and Mg concentrations converge toward seawater values as mixing progresses such that the chemical signature inherited from the water-rock reactions during Step 1 and Step 2 is lost. The K concentrations progress as well toward the seawater value during Step 2, except for degrees of low-temperature basalt alteration > 7% (Figure 4c). In these cases, the fluid mixtures reach equilibrium with K-feldspar at the end of Step 2, the precipitation of which fixes K (Figure S 7 in supplementary information). For Si, the 20 °C fluids retain the chemical signatures of the Step 1 and Step 2 reaction conditions (Figure 4d). The Si concentrations of the fluid mixtures at the end of Step 2 increase with Step 1 temperature maxima. For degrees of low-temperature basalt alteration > 7%, the diversity of Si concentrations in the 20 °C fluids decreases, as observed for pH values, Ca and DIC concentrations.

The Cl and Na concentrations are marginally impacted by the reaction paths (Figure S 8 in supplementary information). Cl is not directly involved in the water-rock reactions and the consumption of H₂O in the formation of hydrated minerals is low enough that the Cl concentration is not impacted significantly. Because of the high Na concentration in seawater, the relative impact of water-rock reactions such as albitization is low and erased by the mixing with seawater during Step 2.

### 3.2.3. Fe and H₂S

As the temperature and rock/water ratio increase during Step 1, the water-rock reactions produce H₂ from the oxidation of ferrous iron in the basalt. Therefore, sulfate is reduced to H₂S, which increases in concentration such that the fluid reaches equilibrium with pyrite (Figure 5a, Figure S 4 in supplementary information). The concentration of Fe increases as well (Figure 5b), controlled by the precipitation of pyrite and Fe-chlorite (Figure S 4 and Figure S 9 in supplementary information).

The redox conditions in the 20 °C fluids at the end of Step 2 are determined by a balance among the reduced species inherited from Step 1, the ferrous iron supplied by the low-temperature basalt alteration during Step 2, and the inputs of oxidizing species during mixing with seawater. The 20 °C fluids produced
by reaction paths with Step 1 temperature maxima > 300 °C and degrees of low-temperature basalt alteration <1% have high concentrations of H₂S (>10⁻⁴ M) (see square-root shaped curve in Figure 5a). Under these reaction conditions, the oxidized species supplied by seawater during mixing are the limiting reactants in the redox reactions taking place during Step 2 because of the large quantity of reduced species inherited from Step 1 including H₂S. In parallel, the Fe concentrations in the 20 °C fluids are low (~10⁻¹² M, Figure 5b) because the high H₂S concentrations maintain equilibrium with pyrite, the precipitation of which fixes Fe.

Several reaction paths produce 20 °C fluids with extremely low H₂S concentrations (<<10⁻¹² M) (Figure 5a). If the Step 1 temperature maximum is ≤300 °C, a lower quantity of H₂S is inherited from Step 1 such that H₂S becomes one of the limiting reactants in the overall redox reactions during Step 2 (see curves dropping vertically during Step 2 in Figure 5a). The reaction paths involving a degree of low-temperature basalt alteration of 1% similarly result in extremely low H₂S concentrations in the 20 °C fluids. This is rather counterintuitive as one could expect more reduced conditions to form due to the supply of ferrous iron from the dissolution of basalt. However, the supply of Fe from low-temperature basalt alteration during mixing results in precipitation of additional pyrite, which fixes a larger amount of H₂S. As a result, the oxidized species supplied by seawater cease to be the limiting reactants in the redox reactions of Step 2. H₂S is effectively completely oxidized at the end of mixing and O₂ accumulates in the fluid mixture (Figure S 10 in supplementary information). The Fe concentrations remain low (<<10⁻¹² M, see curves dropping vertically during Step 2 in Figure 5b) because the 20 °C fluids reach equilibrium with goethite while the O₂ concentration is high (Figure S 11 in supplementary information).

Increasing the degree of low-temperature basalt alteration to 4% re-establishes relatively reduced conditions in the 20 °C fluids yielding H₂S and Fe concentrations of ca. 10⁻⁹ M and 10⁻⁵-10⁻³ M, respectively (Figure 5a and b). The extra supply of Fe from low-temperature basalt alteration consumes the seawater-derived oxidized species during mixing, such that O₂ no longer accumulates in the fluid mixture (Figure S 10 in supplementary information) and, instead, Fe does. As a result, the 20 °C fluids reach equilibrium with
389 both pyrite and goethite at the end of mixing (Figure S 4 and Figure S 11 in supplementary information).
390 For degrees of low-temperature basalt alteration >4%, the Fe concentration in the 20 °C fluids has no degree
391 of freedom because of O₂ depletion, equilibrium with calcite which buffers the pH, and equilibrium with
392 goethite which fixes the Fe. As a result, all reaction paths converge toward a single Fe concentration in the
393 20 °C vent fluids (Figure 5b), as observed for pH, Ca, DIC and Si.

394 4. Discussion

395 4.1. Reaction conditions at Kamaʻehuakanaloa

396 Comparison of the measured temperature and composition of Kamaʻehuakanaloa vent fluids through time
397 (Figure 2) suggests that the overall water-rock interactions that characterize the hydrothermal circulation
398 have remained roughly constant for 30 years and were not substantially modified by the 1996 eruptive event
399 over a period longer than a decade. The 20-50 °C venting regime observed during a series of expeditions
400 since 1987 might, thus, correspond to a chronic state that is only temporarily modified by transient volcano-
401 tectonic events. However, our comparison gives only a broad picture of fluid evolution through time
402 because (1) the morphology of the seafloor hosting venting was completely reshaped by the formation of
403 Pele's Pit during the 1996 eruptive event, (2) not all expeditions sampled the same vent sites, systematically
404 and (3) the datasets for individual samples are not always complete. This large-scale comparison, therefore,
405 does not preclude evolution of vent fluid chemistry and temperature at smaller spatial and temporal scales.
406 Nor do these data allow us to address what happens during the course of a volcanic eruption which, although
407 transient, might release important chemical fluxes that could dominate the time-integrated hydrothermal
408 output from a submarine volcano.

409 Sedwick et al. (1992) used Cl⁻ depletion, alkali metal enrichment, Sr isotopes and correlations between
410 Si and fluid temperature in 5-31°C samples exiting the seafloor at Kamaʻehuakanaloa to infer the occurrence
411 of high-temperature water-rock interactions at depth, as had been done previously for mid-ocean ridge
412 ‘black smokers’ based on the early discoveries at the Galapagos Rift (Edmond et al., 1979). The general
413 model of Si behavior during hydrothermal circulation is that hydrothermal fluids equilibrate with quartz at
high temperature (>300 °C) and then rise so rapidly through the ocean crust that they neither gain nor lose Si, despite the high degree of quartz supersaturation that results from the corresponding pressure drop (Von Damm et al., 1991). However, as mentioned by Sedwick et al. (1992), the low-temperature basalt alteration required to explain the Fe enrichment in Kama‘ehuakanaloa vent fluids could also produce chemical signatures like those expected from the presence of a >300 °C fluid endmember.

In Figure 6, the measured compositions of our Kama‘ehuakanaloa vent-fluid samples are compared to modeled compositions at the end of Step 2, where the fluid mixtures reach temperatures between 50 °C and 20 °C, i.e., the range of temperature measured during the 2018 SUSBEA expedition. Only the results of reaction paths involving mechanical mixing during Step 2 or a narrow degree of low-temperature basalt alteration, *F*, between 2% and 8% are shown here. Further figures showing the full range of reaction conditions considered in our work, across a broader range of *F* values, are provided in supplementary information: Figure S 12 to Figure S 20. Our model shows that while the measured Si concentrations are indeed consistent with mechanical mixing between seawater and a ≥350 °C hydrothermal fluid endmember, the pH and alkalinity values and the concentrations of Ca, K and Fe are not. Rather, the use of multiple tracers allows us to narrow down the range of possible reaction conditions for the Kama‘ehuakanaloa vent fluids and makes it possible to distinguish between the influence of (Step 1) the maximum temperature and the rock/water ratio reached at depth and (Step 2) the degree of low-temperature basalt alteration during fluid upwelling and mixing.

Regardless of Step 1 temperature maxima, mechanical mixing is not able to explain the observed distribution of vent-fluid compositions (e.g., Figure 6, left column). Instead, the modeled Si concentrations and pH values (Figure 6a-d) only migrate towards our measured values as one increases the degree of low-temperature basalt alteration, *F*, during Step 2. For these two parameters, our 2018 data are matched most closely by models in which reaction paths involve either: (Figure 6b) Step 1 temperature maxima of 300-350 °C and a degree of low-temperature basalt alteration of F=2% during Step 2 or (Figure 6c) Step 1 temperature maxima of ≥350 °C and a degree of low-temperature basalt alteration of 4% during Step 2.
Increasing the degree of low-temperature basalt alteration any further only leads to a decrease of the chemical diversity and departure of the modeled fluid compositions from the measured fluid concentrations (Figure 6d and Figure S 12 in supplementary information).

However, if alkalinity is considered (Figure 6e-h), an even narrower range of possible reaction conditions is revealed. The high alkalinity values of at Kamaʻehuakanaloa compared to other submarine vent fields (Table 2) are best explained by reaction paths involving Step 1 temperature maxima of ≥350 °C and a low-temperature basalt alteration of 4% during Step 2 whereas a F value of 2% is less consistent (Figure 6f and g, Figure S 13 in supplementary information). For the case of vent-fluid Ca and K compositions, our approach is less diagnostic: a broad range of reaction paths can account for the measured Ca and K concentrations in the 20-50 °C vent fluids because the modeled concentrations are less sensitive to the reaction conditions than the pH values, Si concentrations or alkalinity (Figure 6i to 6p, Figure S 15 and Figure S 16 in supplementary information). For instance, reaction paths involving Step 1 temperature maxima >300 °C and mechanical mixing during Step 2 or equilibrium mixing with a degree of low-temperature basalt alteration of 2% show some consistency with the measured Ca and K concentrations (Figure 6i, j, m and n) even though those reaction conditions are precluded by the Si and alkalinity data. In contrast, reaction paths with Step 1 temperature maxima of ≥350 °C and a degree of low-temperature basalt alteration of 4% are required to offer consistency with all of our measured pH, Si, alkalinity and Ca concentrations (Figure 6k and o). For higher degrees of low-temperature basalt alteration, F, the diversity of Ca and K concentrations in the 20-50 °C vent fluids decrease, as seen for Si and alkalinity, and the modeled concentrations deviate from our measured vent-fluid values (Figure 6l and p).

Unlike the case for Ca and K, consideration of Fe concentration data *does* help narrow the range of possible reaction conditions at Kamaʻehuakanaloa (Figure 6q to 6t, Figure S 17 in supplementary information). Our measured Fe concentrations are only consistent with a highly constrained set of tested reaction conditions: a Step 1 temperature maximum of ≥350 °C and a degree of low-temperature basalt alteration of F=4% during Step 2 (Figure 6s). Reaction paths involving these same reaction conditions also
predict low H₂ and H₂S concentrations (<10⁻⁹M), which are consistent with those species being below our detection limits during the 2018 field program (Table 1, Figure S 10 in supplementary information). Under these same reaction conditions, the modeled sulfate, Mg and Na concentrations in 20-50 °C vent fluids at the end of Step 2 are also consistent with measured vent-fluid values (Figure S 18, Figure S 19 and Figure S 20 in supplementary information). Note that the measurements of Mg and sulfate concentrations in Kamaʻehuakanaloa vent fluids slightly lower than seawater values suggest that any Mg and sulfate depletion produced by the formation of Mg-smectite and anhydrite during fluid upwelling is not completely overprinted by mixing with seawater in the shallow subsurface (Figure S 1 and Figure S 21).

In addition to the chemical compositions of the vent-fluid samples collected from our expedition, our model is also consistent with mineral observations in high and low-temperature basalt alteration at the seafloor and with mineral observations from Kamaʻehuakanaloa that have been reported in previous studies. As seawater percolates downward beneath the seafloor and is heated above ~150 °C, Mg tends to be fixed into smectite and chlorite (Alt, 1995). At higher temperatures, the precipitation of epidote, amphibole, prehnite and secondary plagioclase controls Ca and Na concentrations (Alt et al., 2010). Consistent with this, our modeled Step 1 hydrothermal fluids reach equilibrium with Ca-Na-saponite, daphnite (Fe-chlorite), epidote, tremolite, prehnite and albite (Figure S 1, Figure S 9, Figure S 2, Figure S 3, Figure S 22, Figure S 23 in supplementary information).

In general, the oxidizing alteration of oceanic basalts at low-temperature in the upper portion of the volcanic section is characterized by the formation of Fe-oxyhydroxides – mostly goethite, replacing olivine – and celadonite and smectite replacing plagioclase and fixing alkalis (Alt, 1995). In extreme cases of oxidizing alteration, K-feldspar replaces plagioclase in the shallowest section of the seafloor (Alt, 1995; Santiago Ramos et al., 2020). Drilling within altered basalts of the Nintoku Seamount, which lies in the Emperor Seamount Chain, revealed a secondary mineralogy made of smectite, Fe-oxyhydroxide, iddingsite and Ca-carbonates, produced by the circulation of large quantities of cold oxidizing seawater (Révillon et al., 2007). Similarly, at Kamaʻehuakanaloa, altered basalts are characterized by extensive hydrothermal
489 deposits of Fe-oxyhydroxides, including goethite, and smectite ranging in composition from Fe-
490 montmorillonite to nontronite (Malahoff et al., 1982; De Carlo et al., 1983; Emerson and Moyer, 2002;
491 Toner et al., 2012; Rouxel et al., 2018). Recently, Ca-carbonates were also reported as ubiquitous secondary
492 minerals at Kamaʻehuakanaloa, filling vesicles of altered basalt at vent sites (Garcia, 2019). Consistent with
493 these observations, reaction paths involving Step 1 temperature maxima of ≥350 °C and a degree of low-
494 temperature basalt alteration of F=4% predict 20-50 °C vent fluids at equilibrium with Na-Ca-Mg
495 saponite/montmorillonite, K-feldspar, goethite, and calcite (Figure S 1, Figure S 7, Figure S 11 in
496 supplementary information and Figure 3d). In contrast to field observations, K-fixation in our model is
497 controlled by the precipitation of muscovite (K-Al-endmember of micas) because celadonite (ferric-mica)
498 is absent from our database. Although nontronite is reported to occur in hydrothermal deposits at
499 Kamaʻehuakanaloa, better consistency between modelled and observed vent-fluid composition can only be
500 achieved when suppressing nontronite precipitation during Step 2. This suggests that either the upwelling
501 fluid in the shallow subsurface does not precipitate nontronite at equilibrium or that the precipitation of
502 nontronite only occurs beyond the end of Step 2, as 20-50°C vent fluids exit the seafloor. Alternatively, the
503 thermodynamic properties of nontronite available in our database may not accurately represent those of
504 nontronite precipitated at Kamaʻehuakanaloa.
505

To summarize, our modeling indicates that a temperature maximum of ≥350 °C and a rock/water ratio
505 of ≥1 during Step 1 together with a degree of low-temperature basalt alteration, F, of ~4% during fluid
506 upwelling and mixing (Step 2) represent the reaction conditions most consistent with production of the
507 vent-fluid compositions observed at Kamaʻehuakanaloa during our expedition. These reaction conditions
508 represent a unifying solution that explains most of the chemical features of the Kamaʻehuakanaloa vent
509 fluids reported here as well as mineral observations that have been reported elsewhere. The general model
510 of simpler mechanical mixing between high temperature fluids and seawater that has previously been used
511 to describe Si behavior during high-temperature hydrothermal circulation at mid-ocean ridges is not
512 applicable at Kamaʻehuakanaloa because low-temperature basalt alteration during fluid upwelling and
513
mixing mobilizes Si, among other species, but can also remove it from solution through the precipitation of alkalic feldspar and smectite.

Other mechanisms beyond our preferred model output might also exist, of course, beyond the range of parameter-space that we have considered. Notably, only a limited set of temperature-rock/water ratio paths during Step 1 were considered. Cool hydrothermal circulation involving very low rock/water ratios (<5·10⁻⁴) at temperature <10 °C (Coogan and Gillis, 2018) was not investigated, but such systems fall below the range of temperatures reported for most vent-fluids exiting the summit of Kamaʻehuakanaloa (Figure 2). Similarly, the effect of conductive heat loss was not investigated in our models. While our mixing model is consistent with both the measured temperature and the measured concentrations of chemical species that behave close to mechanical mixing during Step 2 (Mg, sulfate), this would not be the case if conductive heat loss played a significant role in the temperature decrease of the fluid mixture. The relationship between temperature and the degree of mixing during Step 2 is exponential such that the temperature gradient that drives conductive heat loss decreases exponentially with the degree of mixing. Therefore, it seems reasonable to assign most of the temperature decrease during Step 2 to mixing with seawater.

Like all models, the effectiveness of our model in relating it to the hydrothermal circulation at Kamaʻehuakanaloa is only as good as our assumptions. However, the wide range of reaction conditions considered, and the consistency of the solution presented above in accounting for the array of tracers investigated reassures us of its applicability to Kamaʻehuakanaloa in its current state.

### 4.3. Kamaʻehuakanaloa as a model for low-temperature vent fluids at intraplate volcanoes?

In addition to the maximum temperatures and the rock/water ratios reached at depth at Kamaʻehuakanaloa (Step 1), the low-temperature basalt alteration during fluid upwelling and mixing (Step 2) exerts strong controls on the modeled composition of the 20-50 °C vent fluids. On the one hand, the increase of basalt alteration during Step 2 decreases the diversity of vent-fluid compositions as a function of Step 1 reaction
conditions; on the other hand, it leads to a dramatic change in the dominant reduced species in the 20-50 °C vent fluids: an enrichment in Fe and a concurrent depletion in H₂S and H₂.

The occurrence and extent of low-temperature basalt alteration during fluid upwelling may depend on the geometric arrangement of the steep-sloped Kamaʻehuakanaloa seamount and the associated hydrogeologic plumbing of its hydrothermal convection cell. In parallel work pursued by the SUBSEA team, oxygen and hydrogen isotopic compositions of Kamaʻehuakanaloa vent fluids were used to argue that the seawater recharging the hydrothermal system comes from Pacific Deep Waters at >3000 m, much deeper than the summit crater vents at ~1150-1300m. That study suggests that such a convection cell is made possible specifically because of the steepness of Kamaʻehuakanaloa's flanks which allow deep ocean waters to be drawn in through porous ocean crust across much shorter lateral distances from the site of discharge than would be possible across the more gently-sloping flanks of mid-ocean ridge hydrothermal systems. Simultaneously, this geometric arrangement could accommodate a much longer circulation path from the reaction zone, where maximum temperatures are reached at depth, to the summit where topography forces hydrothermal fluids to vent (Titarenko and McCaig, 2016). The magma chamber at Kamaʻehuakanaloa could be as deep as 7 km below the summit of the volcano (Duennenbier et al., 1997), significantly deeper than at mid ocean ridges. In mid-ocean ridge systems, the hydrogeological plumbing of the hydrothermal convection cell is accommodated by ridge-parallel bounding faults which allow for rapid fluid flow. This contrasts with the Kamaʻehuakanaloa seamount where the stacking of basalt flows can lead to much greater tortuosity in the up-flow path, leading to a much greater surface area of basaltic material being exposed to fluid interactions and, as a consequence, greater basalt reactivity. This combination of greater depth and tortuosity could also lead to significantly longer fluid residence times and/or flow-paths in the subsurface at Kamaʻehuakanaloa when compared to mid-ocean ridge systems. These, in turn, might be expected to give rise to the more extensive low-temperature basalt alteration that we argue for, here, as Step 2 of our conceptual model. If correct, we can hypothesize that active intraplate
seamounts should predominantly produce low-temperature Fe-rich vent fluids, as observed at
Kamaʻehuakanaloa.

Intraplate hydrothermal activity across the vast interior of Earth's Ocean basins has been
disproportionately overlooked, to date, when compared to the attention paid to venting at mid-ocean ridges.
Only four sites of venting on intraplate seamounts have been reported to the InterRidge vents data base
across the entire Pacific (Beaulieu et al., 2013). Interestingly, two of them share characteristics with
Kamaʻehuakanaloa in terms of their reported vent fluid and hydrothermal plume characteristics. In the late
1980s, low-temperature hydrothermal venting was reported at the summit of the Teahitia seamount, the
most westerly active seamount of the Society Islands hotspot in the Central Pacific Ocean (Cheminee et al.,
1989; Binard et al., 1992). Vent fluids were observed exiting from small Fe oxyhydroxide chimneys with
fluid temperatures of ~30 °C, pH values ranging from 5.3 to 5.6, high concentrations of Si (0.8–0.9
mmol/kg), Fe (170–190 µmol/kg) and Mn (6–13 µmol/kg) but no detectable dissolved H₂S (Michard et al.,
1993). More recently, German et al. (2020) measured high concentrations of total dissolvable Fe (≤186
nmol/L) and Mn (≤33 nmol/L) in the hydrothermal plume overlying the Teahitia seamount, comparable to
hydrothermal plume values measured at Kamaʻehuakanaloa (Sakai et al., 1987; Bennett et al., 2011).

At the Vailulu'u seamount, an active submarine volcano located 45 km east of the easternmost island
in the Samoan archipelago, Staudigel et al. (2006) reported low-temperature vents (<81 °C) exiting from
centimeter-to-meter-sized Fe-oxide chimneys together with evidence of high concentrations of dissolved
CO₂. As at Kamaʻehuakanaloa, the Teahitia and Vailulu'u seamounts stand >1500m above the surrounding
seafloor. The similarity between the Kamaʻehuakanaloa, Teahitia and Vailulu'u seamounts suggests that
chronic hydrothermal activity associated with submarine intraplate hotspot volcanoes may predominantly
lead to low-temperature vent fluids enriched in Fe, Mn, Si, CO₂ compared to seawater and depleted in H₂
and H₂S compared to mid-ocean ridge 'black smokers'.

It has long been known that hydrothermal activity can inject detectable amounts of He enriched in ³He
into the oceans from submarine venting, from both mid-ocean ridges and intraplate volcanoes (Lupton et
al., 1977; Lupton, 1998). Recent work from the US GEOTRACES program has shown that important fluxes of Fe can also be injected along with that venting which can be upwelled to the surface ocean (Resing et al., 2015). However, while the principal source for Fe at mid-ocean ridges is considered to be from low-temperature venting with Fe concentrations in the 10-100 µM range (German et al., 2015), the Fe concentrations of vent fluids at intraplate systems like Kama‘ehuakanaloa are 1-2 orders of magnitude higher and are injected into the ocean at much shallower depths beneath the photic zone (Jenkins et al., 2020). Therefore, Fe fluxes associated with intraplate volcanoes may be disproportionately important in terms of global biogeochemical cycles, beyond the deep, dark seafloor.

## 5. Conclusions

New data on submarine vent-fluid compositions acquired at the Kama‘ehuakanaloa seamount during the SUBSEA exploration aboard the *E/V Nautilus* in 2018, coupled with a 30-year record of Kama‘ehuakanaloa fluid composition data, were used to develop geochemical models of water-rock-gas reactions that produce low-temperature vent fluids at intraplate submarine volcanoes. Our model indicates that the compositions of 20-50 °C vent fluids at Kama‘ehuakanaloa, which are enriched in Fe, Si and CO₂ and depleted in H₂S, are not consistent with a simple model of mechanical mixing between seawater and high-temperature vent-fluid. Instead, they appear to reflect an additional stage of subsurface equilibrium mixing between a ≥350 °C hydrothermal fluid end-member and seawater (Step 1), followed by low-temperature chemical alteration of basalt (Step 2) as the resultant hot fluid mixture rises toward the overlying seafloor. Despite the low proportion of additional basalt involved at Step 2 of this process (i.e., only ~4% of the amount of basalt involved in the initial high-temperature reactions at depth), this late reaction step has dramatic impacts on the compositions of the 20-50 °C vent fluids that are emitted into the ocean, decreasing their Si and H⁺ concentrations compared to mechanical mixing and determining which of Fe²⁺ or H₂S+H₂ is the dominant reduced species released from the seafloor. We hypothesize that it is the low-temperature basalt alteration during Step 2 fluid upwelling that gives the Kama‘ehuakanaloa vent fluids their distinctive composition and that this process, in turn, may arise directly from Kama‘ehuakanaloa’s seamount topography.
Specifically, the very steepness of the seamount's flanks could accommodate recharge at great depths where the high-temperature reaction zone could be located, resulting in a much longer circulation pathway from there to the (potentially thousands of meters shallower) summit where the vent-fluids are emitted. Thus, while the recharge path at Kama‘ehuakanaloa may be shorter than for submarine vents at Mid-Ocean Ridges, the up-flow path prior to exit at the seafloor may be significantly longer. Such a circulation geometry could provide an opportunity for extensive low-temperature chemical alteration of basalt via a mechanism that should not be expected at mid-ocean ridges where recharge and discharge are predicted to occur along axis from one another. By contrast, the work we present here could be representative of processes at several other active intraplate seamounts which, as recently shown for Kama‘ehuakanaloa, could contribute substantially to the global Fe-ocean cycle. Such systems could produce low-temperature venting with Fe concentrations 1-2 orders of magnitude higher than at mid-ocean ridges which they would discharge at much shallower depths, injecting this essential micro-nutrient into the oceans closer to the sunlit upper ocean where it can regulate CO₂ draw down by stimulating primary productivity.

## 6. Research data

The output files of the EQ3/6 reaction path calculation can be found in a Mendeley data repository at the following address: https://data.mendeley.com/datasets/n6rbd534h3/1

## 7. Appendix A. Supplementary Material

A supplementary file associated with this paper contains one additional table and 23 figures:

- the chemical composition of the basalt used in the modeling

- the saturation indexes for saponite, epidote, tremolite, pyrite, muscovite, quartz, K-feldspar, daphnite, goethite, anhydrite, prehnite, albite as functions of temperature during modeled reaction pathways

- Na, Cl, O₂ and H₂ concentrations as functions of temperature during modeled reaction pathways
634 - the Si, alkalinity, DIC, Ca, K, Fe, sulfate, Mg and Na concentrations as functions of pH values for
635 modeled and measured 20-50 °C vent fluids

636 ACKNOWLEDGMENTS

637 This work was supported by NASA Planetary Science and Technology Through Analog Research (PSTAR)
638 Program (NNH16ZDA001N-PSTAR) grant (16-PSTAR16_2-0011) to Dr D. S. S. Lim, NOAA Office of
639 Ocean Exploration and Research, Ocean Exploration Trust, NOAA-OER grant NA17OAR0110336, NASA
640 Postdoctoral Fellowship and L'Oréal USA For Women in Science Fellowship to E.T.R. This research used
641 samples and data provided by the Ocean Exploration Trust's Nautilus Exploration Program, Cruise NA100.
642 We thank the cruise expedition Leader N. Raineault, the ROV Hercules/Argus team and the crew of the
643 E/V Nautilus for their assistance while at sea. We are grateful to all members of the SUBSEA team and to
644 the team of the Inner Space Center at the University of Rhode Island Graduate School of Oceanography
645 who made this study possible: John A. Breier, Eric Chan, Josh Chernov, Tamar Cohen, Matthew Deans,
646 Angela Garcia, Michelle Hauer, Renato Kane, Shannon Kobs Nawotniak, David Lees, Justin Lowe, Leigh
647 Marsh, Matthew J. Miller, Miles Saunders, Khaled Sharif, Ashley Shields. We thank the reviewers and the
648 associate editor for helpful and insightful reviews. This is SUBSEA Publication Number: SUBSEA-2018-
649 002.

650
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840

841 Kama‘ehuakanaloa
<table><caption>Table 1. Locations and chemical compositions of fluid samples collected during the 2018 SUBSEA expedition</caption><thead><tr><th>Sample (NA100-)</th><th>002</th><th>001</th><th>039</th><th>040</th><th>019</th><th>054</th><th>053</th><th>061</th><th>060</th><th>080</th><th>079</th><th>059</th></tr><tr><th>Marker<sup>*</sup></th><td>34</td><td>34</td><td>38</td><td>38</td><td>31<sup>g</sup></td><td>2</td><td>2</td><td>57</td><td>57</td><td>57<sup>†</sup></td><td>57<sup>†</sup></td><td>Bottom seawater</td></tr><tr><th>Site<sup>*</sup><br>(lat; long; depth (m))</th><td colspan="5">Spillway Area<br>(18.906; -155.257; 1270±3)</td><td colspan="3">Lohiau area<br>(18.909; -155.257; 1179±2)</td><td colspan="4">Pohaku<br>(18.901; -155.258; 1184±2)</td></tr><tr><th>Temperature max. (°C)<sup>a</sup></th><td>41</td><td>41</td><td>39</td><td>37</td><td>31</td><td>22</td><td>21</td><td>15</td><td>18</td><td>21</td><td>21</td><td>-</td></tr><tr><th>pH (25 °C)</th><td>5.8</td><td>6.0</td><td>6.0</td><td>6.1</td><td>6.0</td><td>6.2</td><td>6.2</td><td>5.4</td><td>5.7</td><td>5.7</td><td>5.7</td><td>7.4</td></tr><tr><th>Alkalinity<sup>b</sup><br>(meq/kg)</th><td>14.3</td><td>14.3</td><td>13.5</td><td>9.2</td><td>8.4</td><td>8.8</td><td>8.9</td><td>6.5</td><td>11.6</td><td>13.6</td><td>12.9</td><td>2.3</td></tr><tr><th>Na (mmol/kg)</th><td>459</td><td>463</td><td>464</td><td>460</td><td>462</td><td>490</td><td>464</td><td>465</td><td>463</td><td>466</td><td>464</td><td>-</td></tr><tr><th>Cl<sup>b</sup> (mmol/kg)</th><td>-</td><td>545</td><td>545</td><td>542</td><td>544</td><td>541</td><td>543</td><td>540</td><td>538</td><td>537</td><td>538</td><td>539</td></tr><tr><th>Mg (mmol/kg)</th><td>51.5</td><td>51.5</td><td>52.0</td><td>52.2</td><td>50.0</td><td>52.5</td><td>52.8</td><td>54.0</td><td>55.0</td><td>54.7</td><td>54.5</td><td>-</td></tr><tr><th>Ca<sup>b</sup> (mmol/kg)</th><td>15.1</td><td>15.8</td><td>15.3</td><td>14.3</td><td>14.9</td><td>16.0</td><td>16.3</td><td>10.9</td><td>11.3</td><td>11.4</td><td>10.6</td><td>10.3</td></tr><tr><th>K (mmol/kg)</th><td>12.2</td><td>-</td><td>12.1</td><td>11.3</td><td>12.4</td><td>16.6</td><td>11.4</td><td>10.7</td><td>11.1</td><td>11.7</td><td>11.7</td><td>-</td></tr><tr><th>Sulfate (mmol/kg)<sup>c</sup></th><td>27.0</td><td>26.7</td><td>26.9</td><td>27.3</td><td>25.8</td><td>27.0</td><td>27.0</td><td>27.9</td><td>27.3</td><td>26.6</td><td>26.8</td><td>28.2</td></tr><tr><th>Si (mmol/kg)</th><td>2.8</td><td>2.7</td><td>2.7</td><td>1.5</td><td>1.9</td><td>1.7</td><td>1.5</td><td>0.6</td><td>1.4</td><td>1.8</td><td>1.9</td><td>-</td></tr><tr><th>Fe (µmol/kg)</th><td>413</td><td>517</td><td>579</td><td>289</td><td>384</td><td>369</td><td>285</td><td>308</td><td>758</td><td>980</td><td>978</td><td>0</td></tr><tr><th>Mn (µmol/kg)</th><td>40</td><td>40</td><td>38</td><td>24</td><td>25</td><td>73</td><td>45</td><td>10</td><td>18</td><td>21</td><td>22</td><td>-</td></tr><tr><th>Al (µmol/kg)</th><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>n.d.</td><td>-</td></tr><tr><th>Sr (µmol/kg)</th><td>104</td><td>103</td><td>103</td><td>95</td><td>98</td><td>-</td><td>102</td><td>87</td><td>88</td><td>87</td><td>87</td><td>-</td></tr><tr><th>Ba (µmol/kg)</th><td>0.7</td><td>0.7</td><td>0.9</td><td>0.7</td><td>0.9</td><td>-</td><td>0.7</td><td>0.7</td><td>1.0</td><td>0.6</td><td>1.0</td><td>-</td></tr><tr><th>B (µmol/kg)</th><td>398</td><td>402</td><td>424</td><td>410</td><td>398</td><td>553</td><td>388</td><td>424</td><td>452</td><td>491</td><td>484</td><td>-</td></tr><tr><th>Li (µmol/kg)</th><td>53</td><td>53</td><td>52</td><td>42</td><td>44</td><td>76</td><td>40</td><td>32</td><td>38</td><td>45</td><td>45</td><td>-</td></tr><tr><th>Rb (µmol/kg)</th><td>5.0</td><td>4.9</td><td>4.7</td><td>3.3</td><td>4.6</td><td>5.8</td><td>3.5</td><td>2.0</td><td>2.8</td><td>3.4</td><td>3.4</td><td>-</td></tr><tr><th>Cs (µmol/kg)</th><td>0.24</td><td>0.24</td><td>0.26</td><td>0.23</td><td>0.26</td><td>0.48</td><td>0.24</td><td>0.24</td><td>0.24</td><td>0.24</td><td>0.25</td><td>-</td></tr><tr><th>Cu (µmol/kg)</th><td>0.9</td><td>1.1</td><td>1.4</td><td>5.3</td><td>1.6</td><td>6.4</td><td>0.5</td><td>1.6</td><td>1.3</td><td>0.9</td><td>1.8</td><td>-</td></tr><tr><th>Zn (µmol/kg)</th><td>27.6</td><td>27.9</td><td>7.7</td><td>6.0</td><td>32.5</td><td>134.4</td><td>12.4</td><td>22.9</td><td>4.3</td><td>4.0</td><td>14.2</td><td>-</td></tr><tr><th>H<sub>2</sub> (µmol/kg)<sup>d</sup></th><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>&lt;3</td><td>-</td></tr><tr><th>CO<sub>2</sub> (mmol/kg)<sup>d</sup></th><td>31.2</td><td>31.2</td><td>30.0</td><td>18.3</td><td>17.4</td><td>14.2</td><td>14.0</td><td>43.8</td><td>94.1</td><td>124.3</td><td>125.8</td><td>-</td></tr><tr><th>CH<sub>4</sub> (µmol/kg)<sup>e</sup></th><td>1.5</td><td>1.3</td><td>1.2</td><td>0.8</td><td>1.4</td><td>0.6</td><td>0.6</td><td>1.6</td><td>2.8</td><td>4.3</td><td>3.8</td><td>-</td></tr><tr><th>H<sub>2</sub>S (µmol/kg)<sup>f</sup></th><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>&lt;1</td><td>-</td></tr></thead></table>

<footnote><sup>a</sup>Marker and site from Rouxel et al. (2018)</footnote>

<footnote><sup>b</sup>Temperature maximum recorded during the IGT sampling</footnote>

<footnote><sup>c</sup>Analysis with ICP-MS, unless otherwise noted: <sup>b</sup>Gran titration, <sup>c</sup>ion chromatography, <sup>d</sup>GC-TCD, <sup>e</sup>GC-FID, <sup>f</sup>spectrophotometry Cline (1969)</footnote>

<footnote><sup>g</sup>Duplicate sample leaked</footnote>

<footnote>n.d. = not detected</footnote>

<footnote>Estimated uncertainties (2σ) are as follows: pH (25°C) values ±0.1 unit; Na, Mg, Cl, Ca, K, sulfate concentrations ±3%; SiO₂, Fe, Mn, Sr, Ba, Li, Rb, Cs, Cu, Zn concentrations ±5%; CH₄ and CO₂ concentrations ±10% and ±5%, respectively.</footnote>
<figure><img src="image_4.png" /><figcaption>Table 1. Vent-fluid compositions from Lō'ihi compared with submarine vent fields from various other geologic settings</figcaption></figure>

<table><thead><tr><th>Vent field</th><th>Seawater</th><th>Lō'ihi <sup>a</sup></th><th>21°N, EPR <sup>b</sup></th><th>Mariner</th><th>Baby Bare</th></tr><tr><th>Geologic setting</th><th>-</th><th>Intraplate volcano</th><th>Mid ocean ridge</th><th>Back arc</th><th>(Spring water)<br>Ridge flank</th></tr></thead><tbody><tr><td>Max. T (°C)</td><td>2</td><td>41</td><td>350</td><td>363</td><td>25</td></tr><tr><td>pH (25 °C)</td><td>7.8</td><td>5.9</td><td>3.4</td><td>2.7</td><td>8.3</td></tr><tr><td>Alk. (meq/kg)</td><td>2.3</td><td>14.3</td><td>-0.4</td><td>-2.25</td><td>0.43</td></tr><tr><td>Cl (mmol/kg)</td><td>546</td><td>546</td><td>489</td><td>605</td><td>554</td></tr><tr><td>SO<sub>4</sub> (mmol/kg)</td><td>28</td><td>26.9</td><td>0.5</td><td>-</td><td>17.8</td></tr><tr><td>Na (mmol/kg)</td><td>468</td><td>460.8</td><td>432</td><td>439</td><td>473</td></tr><tr><td>K (mmol/kg)</td><td>10.2</td><td>12.4</td><td>23</td><td>36.9</td><td>6.9</td></tr><tr><td>Ca (mmol/kg)</td><td>10.3</td><td>15.5</td><td>16</td><td>43.8</td><td>55.2</td></tr><tr><td>Mg (mmol/kg)</td><td>53</td><td>51.5</td><td>0</td><td>1.9</td><td>0.98</td></tr><tr><td>Si (mmol/kg)</td><td>0.1</td><td>2.8</td><td>18</td><td>15.1</td><td>0.4</td></tr><tr><td>Fe (µmol/kg)</td><td>&lt;0.001</td><td>465</td><td>1,664</td><td>12991</td><td>&lt;0.1</td></tr><tr><td>Mn (µmol/kg)</td><td>&lt;0.001</td><td>40</td><td>960</td><td>5723</td><td>7.8</td></tr><tr><td>CH<sub>4</sub> (µmol/kg)</td><td>-</td><td>1.4</td><td>70</td><td>-</td><td>-</td></tr><tr><td>H<sub>2</sub> (mmol/kg)</td><td>-</td><td>b.d.l</td><td>1.7</td><td>-</td><td>-</td></tr><tr><td>CO<sub>2</sub> (mmol/kg)</td><td>-</td><td>31.2</td><td>5.7</td><td>39.8</td><td>0.4</td></tr><tr><td>H<sub>2</sub>S (mmol/kg)</td><td>-</td><td>b.d.l</td><td>7.3</td><td>9.3</td><td>-</td></tr></tbody></table>

<footnote><sup>a</sup> Average values of the 2018 SUBSEA fluids sampled at Marker 34.</footnote>

<footnote><sup>b</sup> Measured composition of vent fluids are corrected for entrainment of seawater during sampling based on the quantitative removal of Mg from the fluid during water-rock reactions.</footnote>

<footnote>Data are from Bruland et al. (2006) for seawater, Humphris and Klein (2017) for 21°N EPR and Mariner vent fields (original data from Von Damm et al. (1985) and Mottl et al. (2011), respectively), Wheat and Mottl (2000) for Baby Bare.</footnote>

<footnote>Abbreviation: Max. T = maximum temperature; Alk. = Alkalinity; - = not determined; b.d.l = below detection limit.</footnote>
Figure 1. (a) Conceptual model of hydrothermal circulation producing low-temperature vent fluids at Kama‘ehuakanaloa. In Step 1, seawater reacts progressively with basalt and volcanic gases (CO₂, SO₂ and H₂O) in an open system along an increase of temperature from 2 °C to a temperature maximum that ranges between 50 °C and 400 °C, and a corresponding increase of rock/water ratio from 0.1 to 1.1. In Step 2, the hydrothermal fluid produced during Step 1 mixes with 2 °C seawater in an open system in which it also continues to react with basalt until the fluid mixture reaches 20 °C. The proportion of basalt and gas reacting during Step 2 is modelled as a fraction, F=0% to F=20%, of the amount of basalt reacting during Step 1. Both mechanical and equilibrium mixing were considered during Step 2. (b) Rock/water weight ratios during Step 1 (black line) and Step 2 (colored lines) are shown as functions of the maximum model temperatures reached during Step 1.

Figure 2. Evolution of Kama‘ehuakanaloa vent fluid compositions and temperatures through time. The blue, orange, white and red boxes represent the periods of pre-eruption, 1988-1993 (Karl et al., 1988; Sedwick et al., 1992, 1994; Hilton et al., 1998), syn-eruption, 1996-1997 (Hilton et al., 1998; Wheat et al., 2000), post-eruption, 2006-2008 (Glazer and Rouxel, 2009; Rouxel et al., 2018) and our SUBSEA data (2018; this study). The numbering above each box indicates the number of measurements for each species within each dataset. Note that no data are available for Na, Cl and DIC concentrations from Glazer and Rouxel (2009) and Rouxel et al. (2018) and that the estimations of DIC concentrations from Karl et al. (1988) and Sedwick et al. (1992) were excluded. The samples considered are from Pele’s vents for the pre-eruption period and from Pele’s Pit (Spillway, Lohiau and Hiolo areas) and the Pohaku area for the post-eruption period. Alk. stands for alkalinity and T (°C) for temperature.

Figure 3. pH values, Ca concentrations, DIC concentrations and calcite saturation indexes as functions of temperature during modeled reaction pathways at Kama‘ehuakanaloa. Seawater composition is shown with an open star and the solid black line represents Step 1 fluid compositions. The color gradation from purple to red represents families of paths for Step 2 fluid compositions originating with
Step 1 fluid compositions calculated for temperatures that range from 50 to 400 °C, at 50°C increments. In each set, the thin dotted line represents mechanical mixing during Step 2. The thickness of the remaining lines increases with the degree of low-temperature basalt alteration, F, during Step 2. For readability, the reaction paths for which F=20% are represented by dashed lines. In subplot (d), the dashed-dotted horizontal black line represents mineral equilibrium.

Figure 4. Sulfate, Mg, K, and Si concentrations as functions of temperature during modeled reaction pathways at Kama‘ehuakanaloa. Seawater composition is shown with an open star and the solid black line represents Step 1 fluid compositions. The color gradation from purple to red represents families of paths for Step 2 fluid compositions originating with Step 1 fluid compositions calculated for temperatures that range from 50 to 400 °C, at 50°C increments. In each set, the thin dotted line represents mechanical mixing during Step 2. The thickness of the remaining lines increases with the degree of low-temperature basalt alteration, F, during Step 2. For readability, the reaction paths for which F=20% are represented by dashed lines.

Figure 5. H₂S and Fe concentrations as functions of temperature during modeled reaction pathways at Kama‘ehuakanaloa. The solid black line represents Step 1 fluid compositions. The color gradation from purple to red represents families of paths for Step 2 fluid compositions originating with Step 1 fluid compositions calculated for temperatures that range from 50 to 400 °C, at 50°C increments. In each set, the thin dotted line represents mechanical mixing during Step 2. The thickness of the remaining lines increases with the degree of low-temperature basalt alteration, F, during Step 2. For readability, the reaction paths for which F=20% are represented by dashed lines.

Figure 6. Si, alkalinity, Ca, K and Fe concentrations as functions of pH values for 20-50 °C vent fluids at Kama‘ehuakanaloa. The colored areas within the frame of each plot represent probability density functions of modeled compositions for vent fluids at the end of Step 2, when fluid mixtures reach
temperature between 50 and 20 °C. Temperature maxima from 50 to 400 °C during Step 1 are shown
with color gradation from purple to red. Open symbols are for pre-eruptive vent-fluid compositions from
Karl et al. (1988) and Sedwick et al. (1992) [squares and triangles, respectively]. The black contour lines
within the frame of each panel indicate the probability density functions for post-eruptive vent-fluid
compositions (Glazer and Rouxel, 2009; Rouxel et al., 2018). Closed circles show the SUBSEA data.
The open star represents background seawater composition. Also shown along the top and right-hand
side of each panel are projections of the same probability density functions onto the corresponding axes
for each plot (a-h) for the SUBSEA data (dark grey) and modeled fluid compositions (colored according
to Step 1 reaction temperature). Reaction conditions during Step 2 arise from mechanical mixing (left
column: a, e, i, m, q) or equilibrium mixing with a degree of low-temperature basalt alteration, F, of 2%,
4% and 8% (2nd, 3rd and 4th column). Note the scale change from the 2nd column (F=2%) to the 3rd
(F=4%) column for the panels showing Fe concentrations. The red box outlines the reaction conditions
that are most consistent with our measured vent-fluid compositions.
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