Unique event plumes from a 2008 eruption on the Northeast Lau Spreading Center

Article

Volume 12, Number 9

16 September 2011

Q0AF02, doi:10.1029/2011GC003725

ISSN: 1525‐2027

Unique event plumes from a 2008 eruption on the NortheastLau Spreading Center

Edward T. BakerPacific Marine Environmental Laboratory, NOAA, 7600 Sand Point Way NE, Seattle, Washington98115, USA ([email protected])

John E. LuptonPacific Marine Environmental Laboratory, NOAA, 2115 SW OSU Drive, Newport,Oregon 97365, USA

Joseph A. ResingJoint Institute for the Study of the Atmosphere and Ocean, University of Washington, Seattle,Washington 98195, USA

Pacific Marine Environmental Laboratory, NOAA, 7600 Sand Point Way NE, Seattle, Washington98115, USA

Tamara BaumbergerDepartment of Earth Sciences, ETH Zurich, Clausiusstrasse 25, CH‐8092 Zurich, Switzerland

Marvin D. LilleySchool of Oceanography, University of Washington, Seattle, Washington 98195, USA

Sharon L. WalkerPacific Marine Environmental Laboratory, NOAA, 7600 Sand Point Way NE, Seattle, Washington98115, USA

Ken H. RubinDepartment of Geology and Geophysics, University of Hawai‘i at Mānoa, 1680 East West Road,Honolulu, Hawaii 96822, USA

[1] The creation of ocean crust by lava eruptions is a fundamental Earth process, involving immediateand immense transfers of heat and chemicals from crust to ocean. This transfer creates event plumes(“megaplumes”), massive ellipsoidal eddies with distinctive and consistent chemical signatures. Here wereport the discovery of unique event plumes associated with a 2008 eruption on the Northeast Lau SpreadingCenter. Instead of a large plume hundreds of meters thick, we detected at least eight individual plumes, each∼50m thick and apparently only 1–3 km in diameter, yet still rising 200–1000m above the eruption site. Lowand uniform 3He/heat (0.041 × 10−17mol/J) and dissolvedMn/heat (0.04 nmol/J) ratios in water samples werediagnostic of event plumes. High H2 concentrations (up to 9123 nM) and basalt shards confirmed extensiveinteractions between molten lava and event plume source fluids. Remote vehicle observations in 2009mapped a new, small (1.5–5.8 × 106 m3) lava flow. Our results suggest that event plumes are more variable,and thus perhaps more common, than previously recognized. Small event plumes may be preferentiallyassociated with small or sheet‐flow eruptions, and massive event plumes with slowly extruding pillowmounds 25–75 m thick. Despite this correlation, and high H2 concentrations, existing theory and seafloorobservations argue that cooling lava cannot transfer heat fast enough to create the buoyancy flux required

Copyright 2011 by the American Geophysical Union 1 of 21

http://dx.doi.org/10.1029/2011GC003725

for event plumes. The creation of event plumes under a broad range of eruption conditions provides newconstraints for any theory of their formation.

Components: 12,000 words, 12 figures, 2 tables.

Keywords: Lau Basin; event plumes; hydrothermal venting; seafloor eruption.

Index Terms: 3001 Marine Geology and Geophysics: Back-arc basin processes; 3017 Marine Geology and Geophysics:Hydrothermal systems (0450, 1034, 3616, 4832, 8135, 8424); 3075 Marine Geology and Geophysics: Submarine tectonicsand volcanism.

Received 24 May 2011; Revised 27 July 2011; Accepted 7 August 2011; Published 16 September 2011.

Baker, E. T., J. E. Lupton, J. A. Resing, T. Baumberger, M. D. Lilley, S. L. Walker, and K. H. Rubin (2011), Unique eventplumes from a 2008 eruption on the Northeast Lau Spreading Center, Geochem. Geophys. Geosyst., 12, Q0AF02,doi:10.1029/2011GC003725.

Theme: Assessing Magmatic, Neovolcanic, Hydrothermal, and Biological Processesalong Intra-Oceanic Arcs and Back-Arcs

Guest Editors: J. Resing, K. Rubin, T. Shank, W. Chadwick, and D. Butterfield

1. Introduction

[2] Intrusions and eruptions of magma build theocean crust, yet the physical and chemical interac-tions that occur between crust and ocean during theseevents remain poorly understood. The serendipi-tous discovery of an event plume, or “megaplume,”on the Juan de Fuca Ridge in 1986 was the firstindication that eruptions could be accompanied bybrief and massive releases of hot, chemical‐richfluids [Baker et al., 1987]. The 1986 event plume, aoblate ellipsoid 700 m thick and 20 km in horizontaldiameter (∼150 km3) with a total hydrothermalheat anomaly of ∼1017 J [Baker et al., 1989], wasapparently accompanied by an eruption of ∼0.05 km3

of lava [Chadwick and Embley, 1994; Chadwicket al., 1991]. Since then, three additional ridgecrest eruption events on the Juan de Fuca and Gordaridges have produced six separate event plumes(“Confirmed” in Table 1) with volumes rangingfrom ∼15 to 128 km3 [Baker, 1998]. In addition,five plumes (“Possible” in Table 1) with somecharacteristics similar to event plumes have beenobserved at ridges worldwide, with spreading ratesfrom 11 (Gakkel Ridge) to 140 (Manus Basin) mm/yr (Figure 1). None of the plumes outside the NEPacific, however, were mapped in three dimensionsor could be linked to a specific seafloor eruption.

[3] The most fundamental characteristics of allevent plumes to date are their large size and chem-ical uniformity. Confirmed diameters range from5 to 20 km (and perhaps up to 70 km [Murton et al.,2006]), and thicknesses from 0.5 to 1.2 km. Con-

centrations (in terms of the species/heat ratio) ofdiagnostic magmatic tracers such as 3He and dis-solved Mn (DMn) are nearly uniform among allevent plumes and lower than the same ratios foundin chronic (non‐event) plumes and high‐temperatureseafloor discharge [Lupton et al., 1989, 1999a,2000; Massoth et al., 1995, 1998]. Thick, high‐rising, and broadly dispersed plumes can also occurwithin very deep axial valleys, such as the GakkelRidge, where vertical density gradients are low andplumes become trapped by unbroken valley walls[e.g., Edmonds et al., 2003]. However, these plumeshave neither the three‐dimensional symmetry norchemical characteristics of event plumes.

[4] The physical and chemical uniformity of allevent plumes argues for a common origin, butno conclusive formation theory has yet emerged.Confirmed event plumes in the northeast Pacificoccurred immediately following seismicity indica-tive of magma migration and crustal rifting [Dziaket al., 2007], and coincident with seafloor lavaeruptions [Chadwick and Embley, 1994; Chadwicket al., 1995, 1998]. Consequently, proposed for-mation hypotheses include the release of pre‐formedhydrothermal fluids during crustal rupturing [Bakeret al., 1989; Cann and Strens, 1989; Lupton et al.,1999a, 2000; Wilcock, 1997], the rapid cooling ofan intruded dike [Lowell and Germanovich, 1995],and the rapid cooling of an erupting lava flow[Butterfield et al., 1997; Palmer and Ernst, 1998,2000; Clague et al., 2009].

[5] Here we report observations of hydrothermaldischarge associated with a seafloor eruption in

GeochemistryGeophysicsGeosystems G3G3 BAKER ET AL.: UNIQUE EVENT PLUMES AT A RIDGE ERUPTION 10.1029/2011GC003725

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November 2008 on the Northeast Lau SpreadingCenter (NELSC), a back‐arc ridge behind the Tongaarc (Figure 2). In May 2009 a response cruisesponsored by NOAA VENTS and NSF [Resing andEmbley, 2009] made repeat water column mea-surements at the presumed eruption site, and usedthe remotely operated vehicle (ROV) JASON toconfirm that lava was erupted in 2008, collect rocksamples, and sample vent fluids. The eruption mayhave been one of a series of small fissure eruptionsat this site that began as early as mid‐2008 [Rubinet al., 2009]. Associated with the eruption was aunique suite of small plumes matching the distinc-tive chemical composition of previous event plumes,but with volumes 10–1000 times smaller. This rel-atively small event extends the correlation betweenevent plume and lava heat content across two ordersof magnitude. The discovery of extraordinarily highconcentrations of H2 gas and abundant glass shardsin the plumes confirms the interaction of event

plume fluids and molten lava [Kelley et al., 1998].Despite these observations, a detailed analysis oflava cooling rate using existing theoretical calcula-tions and field observations shows that heat transferfrom cooling lava rapid enough to create the buoy-ancy flux needed to generate event plumes (hours,not days) is highly improbable. Our observationsprovide new constraints for theories of why seaflooreruptions are associated with brief and massivedischarges of unique hydrothermal fluids.

2. Geological Setting

[6] The NELSC is one of several spreading ridgesaccommodating back‐arc extension behind theTonga arc [Hawkins, 1995; Pelletier et al., 1998;Zellmer and Taylor, 2001] (Figure 2). In the north-ern Lau Basin, the ridges forming the MangatoloTriple Junction and the Fonualei Ridge SpreadingCenter have opening rates of 85 to 94 mm/yr

Table 1. Physical Characteristics of Event Plumes and Associated Eruptions

Event Location

SpreadingRate

(mm/yr)

HorizontalDiameter(km)

Thickness(km)

Heata

(1016 J)Lava Heatb

(1016 J) References

ConfirmedEP86 Juan de Fuca Ridge 55 20 0.7 19.3 19.4–25.3 Baker et al. [1987];

Chadwick et al. [1991]EP87A JDFR 55 16 0.6 6.9 NDc Baker et al. [1989]EP93A JDFR 55 6.5 0.6 0.89EP93B JDFR 55 9 0.7 2.2EP93C JDFR 55 13 0.7 3.9Total 93A‐C 7.0 4.3 Baker et al. [1995];

Chadwick et al. [1998]EP96A Gorda Ridge 55 14 1.2 5.1EP96B Gorda Ridge 55 12 0.8 1.2Total 96A‐B 6.3 8.9–13.9 Baker [1998];

Chadwick et al. [1998]EP08A NELSC ∼90 2.2 0.05 0.036EP08B NELSC ∼90 3.2 0.05 0.038EP08C NELSC ∼90 2.2 0.05 0.016EP08D NELSC ∼90 2.3 0.05 0.013EP08E NELSC ∼90 1.1 0.04 0.0028EP08F NELSC ∼90 1.3 0.065 0.0053EP08G NELSC ∼90 1.8 0.05 0.0034EP08H NELSC ∼90 1.8 0.07 0.0095Total 08A‐H 0.12 0.2–1.2 This paper

PossibleEP87B N Fiji Basin 72 ND 0.6 ND ND Nojiri et al. [1989]EP90 Manus Basin ∼140? ∼70 0.2 ND ND Gamo et al. [1993]EP98 JDFR 55 >15? 0.3 ND 8.9–36.9 Baker et al. [1999];

Embley et al. [1999]EP01 Gakkel Ridge 11 >9? 1.2 ND ND Edmonds et al. [2003]EP03 Carlsberg R. 30 >70 0.7 ND ND Murton et al. [2006]

aWhole plume inventory, corrected for stratification effects on D�.bAssuming total lava heat = latent heat + specific heat of melt cooling (1200° to 1000°C) + specific heat of basalt cooling (1000° to 0°C) =

(lava volume m3)(2700 kg/m3)(4 × 105 J/kg) + (200°C)(1200 J kg−1°C−1) + (1000°C)(1200 J kg−1°C−1).cND, not determined.


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[Zellmer and Taylor, 2001], implying a similar ratefor the NELSC. The southernmost segment of theNELSC is a knife‐edge ridge ∼15 km long with abathymetric high in the middle and prominent cones(Maka and Tafu) at either end (Figure 2). In 2008Nautilus Minerals Inc. made the first visual confir-mation of high‐temperature venting (∼315°C) onthe summit of Maka, the southern of these two cones(J. J. Lowe, Nautilus Minerals Inc., personal com-munication, 2008).

3. Methods

[7] Our main survey tool was a Sea Bird 911plusConductivity‐Temperature‐Depth (CTD) package,with light backscattering and oxidation‐reductionpotential (ORP, sometimes referred to as Eh) sensors.A total of 15 vertical casts and tow‐yos were con-ducted on and around the southern segment of theNELSC during 20–27 November 2008 (Figure 2).The voltage output of the light‐backscattering sen-sors is equivalent to nephelometric turbidity units(NTU) [American Public Health Association, 1985];DNTU is the value in ambient nonplume water.ORP is very sensitive to short‐lived reduced che-micals in hydrothermal plumes, such as Fe+2 andHS‐ [Walker et al., 2007]. TheORP sensor measuresthe electrode potential (E (mV)) between seawaterand a reference solution. Absolute values of E canvary because of instrumental drift and hysteresis (theresponse is instantaneous but recovery time can lastseveral to tens of minutes), especially in concen-

trated plumes, so we rely on the time derivative,dE/dt (mV/s) [Nakamura et al., 2000] to identify theprecise location of anomalies. Because E declineswhen it encounters reduced substances, the anoma-lies are negative. Miniature Autonomous PlumeRecorders (MAPRs) [e.g., Baker et al., 2010], withidentical scattering and ORP sensors, supplementedthe CTD on some tows.

[8] We calculate the hydrothermal temperatureanomaly (D�) in a neutrally buoyant plume fromthe expression

D� ¼ �� m0 �m1�� m2�2�

� �

where � is potential temperature, s� is potentialdensity, and m0, m1, and m2 are constants in alinear or polynomial regression between � and s� inhydrothermally unaffected water around the plume.Because we will compare species/heat ratios inplumes in the Lau Basin with those in the northeastPacific, we use the technique of McDougall [1990]and McDuff [1995] to adjust for the effect ofhydrography and vent fluid characteristics on theobserved D�. Those authors show that the plumeheat flux at the equilibrium level, D�Q, where Q isthe fluid volume flux, equals the vent heat flux,D�vQv, when multiplied by two correction terms,one related to ocean stratification and one to thevent source �v and salinity (Sv). Thus

D�Q ¼ D�v=Qv½ � 1= 1� Rð Þ½ � 1� R=Rvð Þ½ �

Figure 1. Locations of confirmed (black labels) and suspected (gray labels with a “?”) event plumes. The numberafter “EP” gives the discovery year, and the letters indicate successive event plumes in the same year. Black lines arespreading ridges, blue lines are trenches.


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where R = [(a/b)(d�/dS)], Rv = [(a/b)(d�v/dSv)],d�/dS is the vertical �‐S gradient of the ambientwater column through which the plume rises,d�v/dSv is the ratio of the differences in � and S

between seawater and vent fluid, a is the coeffi-cient of thermal expansion, and b is the coefficientof haline contraction (see the Notation list at theend of the paper). To ensure that the plume and

Figure 2. Bathymetry of the southern segment of the Northeastern Lau Spreading Center (NELSC), showing verticalCTD casts from 2008 (white dots), MAPR profiles from 2004 (blue circles) [German et al., 2006], CTD tows from2008 (colored straight lines), and CTD cast 11 and tow 12 from 2006 (red dot and line) [Kim et al., 2009]. Inset showsthe boundaries separating the Australian (A), Tongan (T), and Niuafo’ou (N) plates [Zellmer and Taylor, 2001]. Plateboundary acronyms in addition to the NELSC include: NWLSC, Northwest Lau SC; PR, Peggy Ridge transformfault; LETZ, Lau Extensional Transform Zone; CLSC, Central Lau SC; ELSC, Eastern Lau SC; FRSC, FonualeiRidge SC; and MTJ, Mangatolu Triple Junction. Double‐headed arrows indicate full spreading rates at ridges; single‐headed arrows give GPS velocities of T relative to A.


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vent heat fluxes agree, the observed D� must bemultiplied by a correction factor, ([1/(1 − R)][1 −(R/Rv)])

−1.

[9] Because we have no information on vent sourceS during the creation of any event plume, weassume, for consistency with other event plumes,that Sv = S and set Rv = ∞. For the northeast Pacific,R = −0.97, giving a correction factor of 0.51 andindicating that the observed D� values must bemultiplied by a correction factor of 2. For the LauBasin, R = −2.4, the correction term is 0.29, and theobserved D� values must be multiplied by a cor-rection factor of 3.4. Lavelle et al. [1998] calculatedsimilar correction factors using a more generalizedapproach that extended the corrections to nonlinearhydrographic profiles and event plumes. They alsoshowed that the corrections are independent of theentrainment coefficient of the ascending plume.

[10] Although we have no information to makecorrections for Sv, it is instructive to understand theeffect such changes would have on the D� calcu-lations (Figure 3). In the northeast Pacific, theeffect is not drastic. For vent fluid temperaturesof 350° and 700°C a significant change in thecorrection factor (>2 times increase) occurs onlyfor 350°C fluids with salinities fresher than ∼5 psu.At the NELSC, a steeper d�/dS gradient in the localwater column leads to much greater variability. At350°C, Rv is either greater than or less than Rdepending on dSv. Around the dSv value of −24.64

(i.e., a vent fluid S of ∼10 psu), the correction factortakes on very high positive or negative values,leading to high variability in the corrected D�. At700°C, the correction factor stays monotonic butincreases steadily as Sv decreases. The high sensi-tivity of the correction factor to low vent fluidsalinity provides a strong constraint on possiblevalues of Sv for the source fluids of the NELSCplumes.

[11] Water samples were collected on every CTDcast using PVC sampling bottles. Samples forhelium isotopes were sealed into copper tubingusing a special hydraulic crimping device [Youngand Lupton, 1983]. Helium isotope ratios weremeasured at the NOAA/PMEL laboratory in New-port, Oregon, using a dual collector mass spectrom-eter designed specifically for helium measurements[Lupton, 1990]. We determined the regional 3Heconcentration profile from background stationsoccupied away from the eruption site during thecruise and then subtracted a depth‐appropriatebackground value from all plume samples. Imme-diately after sample collection, dissolved H2 wasmeasured shipboard using a standard headspaceequilibration technique as described by Kelley et al.[1998]. Dissolved Mn was determined with a pre-cision of ∼1 nM by the direct injection method asdescribed by Resing et al. [2007].

[12] We can evaluate the accuracy of correctedspecies/heat ratios in plumes by comparing them

Figure 3. The effect of vent fluid salinity on the magnitude and sign of theD� correction factor for high‐temperaturevent fluids in the Lau Basin and northeast Pacific regions. For positive D� anomalies observed in the Lau Basin,∼350°C source fluids cannot be fresher than ∼10 psu (chlorinity ∼230 mmol/kg).


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to values measured on samples collected from ventorifices located immediately below the plume sam-ples. Regressing D� against 3He for eight plumesamples <300 m above the summit of Maka (castV08C‐27) gave a 3He/heat ratio of 3.3 × 10−17 mol/J(r2 = 0.87, ± 0.5 × 10−17 mol/J at the 95% C.I.), inreasonable agreementwith the value of 5 × 10−17mol/Jfrom high‐temperature (315°C) fluids on Makameasured on samples collected in 2009. This source‐plume agreement demonstrates that we can confi-dently distinguish the order‐of‐magnitude differencesin 3He/heat ratios found between chronic (non‐eruption) plumes and those generated by eruptionevents (as discussed below).

4. Plume Observations and Comparisonto Previous Event Plumes

4.1. Hydrothermal Setting

[13] The first exploration for hydrothermal activityon the NELSC was a series of widely spaced light‐scattering profiles collected in 2004, of which onlythree were over the southern segment [German et al.,2006] (Figure 4). The profiles over Maka and thesegment center show a broad plume layer between1200 and 1500 m (a weak plume between 1250 and1350 m was also present on a profile ∼6 km south-west of Maka). No anomaly was seen over Tafu, thenorthern cone. In 2006, Kim et al. [2009] conducteda tow (CTD 12) and cast (CTD 11) along thesegment, again finding a thick plume over Maka.

No shallow plume was seen between the cones asin 2004, but weak, near‐bottom, light‐transmissionanomalies occurred near the segment center at∼15.395°S (Figure 4). Kim et al. [2009] also detec-ted a thin plume apparently originating from thesummit of Tafu, and dredged a high‐temperaturechimney at ∼1900 m on the northwest flank ofMaka, coincident with a water column anomalyinterpreted as a buoyant plume.

4.2. Physical Characteristics

[14] Previous event plumes have differed fromchronic plumes in their symmetry and size, indic-ative of their origin by brief and massive injectionsof hot fluids. The shape of confirmed event plumesapproximates an oblate ellipsoid, with diametersranging from 5 to 20 km and thicknesses from 500to 1200 m (Table 1). Event plumes detected overthe NELSC were unique in both their small sizeand their abundance.

[15] In 2008, high‐rising plumes were initiallydetected on tow‐yo T08C‐07 (20 November), ourfirst operation at the NELSC (Figure 5). On thefinal upcast, a series of thin plumes with highDNTU and D� values was detected between 600and 1200 m, as much as 1000 m above the adjacentridge crest axis. Extreme dE/dt anomalies, <−6mV/s,confirmed that the plumes were relatively young(also confirmed by extraordinarily high H2 con-centrations; see section 4.3.2). Our next tow alongthe ridge, T08C‐09 (21 November), found another

Figure 4. Bathymetric profile along the NELSC beneath CTD 12 (gray sawtooth line) of Kim et al. [2009] (seeFigure 2). Actual summit height of Maka shown in gray. Plume distribution from percent light transmission mea-surements on CTD 12 shown in gray contours. MAPR profiles from 2004 [German et al., 2006] shown in red; verticalaxis gives full depth of each profile, horizontal axis is DNTU. Vertical black lines indicate other CTD cast locationsfrom Kim et al. [2009].


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remarkable series of thin and intense plumesbetween 1600 and 900 m (Figure 6). We interpretthis distribution as a record of sequential and discretefluid discharges creating plumes of diminishingbuoyancy flux (i.e., diminishing rise height) injectedfrom the ridge crest into a southward current. (Notethe southward advection of the plume from Maka inFigure 5.) Maximum dE/dt anomalies were muchsmaller than on T08C‐07, implying somewhat olderplumes. Plumes shallower than ∼1500 m had aconsistent thickness of ∼40–70 m. Maximum riseheight and D� values were comparable to the orig-inal 1986 “megaplume,” EP86 [Baker et al., 1987],although for that event a single plume encompassedthe entire water column from 1200 to 2000 m.

[16] This plume distribution lasted only a few days.By the time of the final CTD profiles of the cruise(V08C‐18 (24November), V08C‐27 (27November),T08C‐18 (27 November)), we detected no plumesshallower than those from theMaka summit (∼1300mminimum depth), and even the Maka plumes wereabsent from profiles over the eruption site. This rapidadvection of the event plumes away from the erup-tion site also implies that the plumes we observed on20 and 21 November were no older than approxi-mately one week.

[17] Using the conservative assumptions that theplumes were symmetrical and that T08C‐09 passedthrough the center of each (its path followed

Figure 5. T08C‐07 transect showing (a) dE/dt anomalies, (b)DNTU contours and CTD tow‐yo (thin black line), and(c) map of the ridge crest showing the CTD tow path (white line) and ridge axis (dashedwhite line). In Figure 5b, the graycross‐section indicates the bottom profile along the track line; white dashed line above shows the ridge crest profile(from Figure 5c) projected onto the tow track profile. Small diamonds show sample locations along the CTDtow‐yo. Event plumes were first observed at the end of this tow as a series of thin plumes with intense anomalies.The extreme dE/dt anomalies of the event plumes emphasize their young age relative to the Maka plume.


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the axis of the ridge crest), the total volume of allthe event plume lenses seen on that tow is ∼1 km3,∼15 times less than EP93A, the smallest eventplume previously recorded [Baker et al., 1995].(For operational reasons, no cross‐axis tows wereconducted to confirm this assumption.) Such thinplumes are not without precedent during eruptionevents, however. A CTD cast at the CoAxial erup-tion site, serendipitously conducted while the seis-mic activity was still underway, found a single50 m thick plume layer ∼600 m above the seafloor[Baker et al., 1995]. A vertical profile throughEP93A, about 10 d later, also showed pronounced

layering throughout its 600m thickness. Thus smallereruptions associated with smaller event plumes (e.g.,at the NELSC) may release insufficient energyto uniformly mix the water column over verticallength scales of hundreds of meters. The small,layered plumes we observed in 2008 also suggestthat event plume structure grades from a large,highly symmetrical vortex (i.e., a “megaplume”)created during a large eruption, such as the 1986or 1987 Juan de Fuca Ridge events, to an array ofthin, layered “blobs” as created during the NELSCeruption. These blobs could be symptomatic ofpulsed eruption activity.

Figure 6. T08C‐09 transect showing (a) dE/dt anomalies, (b) DNTU contours and CTD plus MAPRs tow‐yo (thinblack lines), and (c) map of the ridge crest showing the CTD tow path (white line) and ridge axis (dashed white line).In Figure 6b, the gray cross‐section indicates the bottom profile along the track line; white dashed line above showsthe ridge crest profile (from Figure 6c) projected onto the tow track profile.DNTU values were high but dE/dt anoma-lies were much weaker than on T08C‐07. Plumes appeared to originate from the area of new lava (Puipui) mapped inMay 2009 (red area on profile and on T9 track in Figure 6c). H2/heat (nmol/J) ratios shown next to each bottle sample(diamonds) (including the one sample (T7, white dot) from tow T08C‐07 shown at its depth and location). Blackdashed ellipses mark the cores of eight separate event plumes used to calculate total event plume volume. Inset showsthe size and location of a typical basalt shard suspended in the plumes.


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[18] Because our physical and chemical plumemeasurements strongly suggested a very recenteruption, perhaps only days or hours old, we had theopportunity to characterize the plumes within ∼50 mof the seafloor at a remarkably early stage after aneruption. In 2008 our sole clue to the eruptionlocation was the intersection of the seafloor andthe plumes seen on tow T08C‐09 (Figure 6). AD� of 0.7°C recorded just above the seafloor at15.389°S during T08C‐09 was convincing evi-dence of the source location, so we conducted twoadditional tows along the ridge crest that wouldpass through that area 18 h (T08C‐10) and 7 days(T08C‐18) after T08C‐09 (Figure 7). During eachtow, the CTD altitude was kept at <50 m (andusually <30 m) through the target area. A ROVtransect conducted in July 2008 along this part ofthe ridge crest had found only the low‐temperatureNautilus vent (J. J. Lowe, Nautilus Minerals Inc.,personal communication, 2008).

[19] Temperature and dE/dt anomalies along thosethree tows suggested the presence of numerous fluidsources concentrated along the shallowest portion ofthe ridge crest (<∼1650 m), between 15.381° and15.392°S (Figure 7). D� values >0.2°C and dE/dtanomalies <−0.5 mV/s were common. The tows

also identified an apparent low‐temperature ventsite with a strong dE/dt response south of 15.40°S,deeper than ∼1800 m, that apparently correspondsto the deep plume seen between 15.4° and 15.41°Son T08C‐09 (Figure 6).

[20] A repeat CTD tow conducted in May 2009found no evidence of continuing discharge fromthe shallowest portion of the ridge crest. However,detailed mapping of the eruption area by JASONdid confirm the presence of fresh, unsedimentedlava, called the Puipui flow, extending about 1.7 kmalong the ridge crest, with a cross‐axis width of250–400 m [Rubin et al., 2009]. Over most of itslength, the flow comprises highly vesicular (up to∼50% near the eruptive vents) basalts in the formof high‐effusion‐rate sheets and lobate flows, withpillow lavas mostly limited to flow margins [Clagueet al., 2010]. JASON also observed weak diffuseventing (<20°C) at two locations on the lava flow,including the Nautilus site discovered in 2008(Figure 7).

[21] The high D� anomalies seen both in the eventplumes and just above the seafloor over the lavaflow provide an important constraint on the salinityof the source fluids. For the local hydrography, the

Figure 7. Hydrothermal temperature (D�) and ORP (dE/dt) anomalies in the near‐bottom water over the lavaemplacement zone. Grey patch on the axis defines the approximate known extent of the 2008 Puipui flow; cyan starsshow the Nautilus diffuse vent (to the north) and a second diffuse vent discovered in 2009. Data and tow paths ofT08C‐09 (red), −10 (blue), and −18 (green) appear only where the CTD altitude was <50 m. D� scale is shownto the top right, dE/dt scale to the bottom left.


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D� correction factor becomes unreasonably highor even unstable as high‐temperature source fluidsbecome fresher (Figure 3). Thus it is probable thatthe source fluids for both sets of plumes could nothave been of unusually low salinity. This is a sig-nificant restriction, since immediate post‐eruptionfluids from high‐temperature vents can have Sv‐Svalues <−30 [Von Damm et al., 1995; Lilley et al.,2003], which would result in strongly negative D�anomalies at the NELSC.

4.3. Chemical Characteristics

4.3.1. 3He and DMn

[22] The most diagnostic characteristics of eventplume water chemistry are the uniquely low andconsistent ratios of 3He and DMn to hydrothermalheat (Figures 8, 9, and 10 and Table 2) [Lupton et al.,1999a, 2000;Massoth et al., 1995, 1998]. Historicalmeans ( ±1s) are 0.18 ± 0.022 × 10−17 mol/J for 3He/heat and 0.059 ± 0.014 nmol/J for DMn/heat. (Note

that both values are lower than previously reportedbecause of the hydrographic temperature correc-tions to plume D�, as discussed in section 3. SomeDMn values reported by Massoth et al. [1995,1998] are actually total dissolvable Mn, which canbe a few percentage points higher than true DMn.)Because both 3He and hydrothermal heat are con-servative tracers, we focus on their ratio to placeevent plumes in the context of magmatic activity.

[23] Lupton et al. [1999a, 2000] noted that 3He/heatratios fall into two groups: low ratios (mostly <1 ×10−17 mol/J) in event plumes and high‐temperaturefluids from magmatically quiescent locations, andhigh ratios (mostly >1 × 10−17 mol/J) from eruption‐associated chronic plumes, especially plumessampled very close to newly erupted lava flows.The two groups could thus be considered end‐members that, when combined, are consistent withthe theoretical 3He/heat ratio within the upper mantle,∼2 × 10−17 mol/J [Lupton et al., 1989].

Figure 8. Scatterplots of (a) 3He/heat (red symbols) and DMn/heat (blue symbols) from tows T08C‐07, −09, −10, and−18, and (b) H2/heat (green triangles) ratios from tows T08C‐07 and −09. In Figure 8a gray curves show plume distri-bution (D�) from T08C‐09. Open symbols show the ratios for samples in the plume from the summit of Maka, whichwere similar to values from the deep plumes on the ridge. Means and standard deviation for samples <1400m (excludingthe Maka plume) are 7.3 ± 2.9 × 10−19 mol/J (n = 11) for 3He/heat, 0.06 ± 0.025 nmol/J (n = 11) for DMn/heat, and 3.6 ±1.7 nmol/J (n = 10) for H2/heat. Means and standard deviation for all other samples are 1.5 ± 1.7 × 10−17 mol/J (n = 38)for 3He/heat, 0.46 ± 0.29 nmol/J (n = 32) for DMn/heat, and 0.46 ± 0.60 nmol/J (n = 6) for H2/heat.


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[24] A compilation of time series measurementsin chronic plumes following an eruption shows thatthose 3He/heat ratios, unlike in event plumes,vary not only temporally but also among eruptions(Figure 11). 3He/heat ratios in post‐eruption chronicplumes can exceed the magmatically quiescentvalue typical of each eruption site by a factor oftwo or three. Lilley et al. [2003] observed a similartrend in post‐eruption time series measurements ofvent fluids following the 1991 eruption near 9.9°Non the East Pacific Rise and the 1999 magmaticevent at the Endeavor segment of the Juan de FucaRidge. 3He/heat ratios in post‐eruption chronicplumes can remain high for at least a year. High‐resolution sampling suggests that the highest valuesoften occur some months after the eruption, per-haps when the extruded and/or intruded magma hascooled enough that cracks allow seawater accessto 3He and other volatiles in the basalt [e.g., Bakeret al., 1995; Lupton et al., 1999b]. The high vari-ability of post‐eruption chronic plumes indicates

that their source fluids have an origin separate fromthose of event plumes.

[25] Ratios found in the 2008 event plumes andassociated discharge had many similarities to pre-vious eruptions, but with some distinct differencesthat make the 2008 Lau plumes unique (Figures 9,10, and 11 and Table 2). For 3He/heat, event plumesamples were the lowest ever measured, 0.041 ×10−17 mol/J. As at all other eruption sites, the post‐eruption plumes over the lava and <∼100 m abovethe seafloor (i.e., 1500–1650 m) had ratios (0.1 to0.3 × 10−17 mol/J) consistently higher than the eventplumes; these ratios showed no temporal trend onfive casts collected over a week’s time. However,these plumes were unique among other eruptions inthat their 3He/heat ratios were 10–20 times lowerthan historical chronic plume and vent samplesassociated with recent magmatic activity on othermid‐ocean ridges. And unlike other eruptions, thesepost‐eruption plumes had 3He/heat ratios muchlower than the non‐eruption‐associated dischargesources on the NELSC. Plumes sampled in 2008

Figure 9. 3He/heat ratios (mol/J) from the 2008NELSCeruption site for samples in event plumes (red dots, y =2.7 × 10−16 + 4.1 × 10−19x, r2 = 0.68), over the lava flowat depths of 1500–1650m (blue squares, y = 1.1 × 10−16 +2.1 × 10−18x, r2 = 0.79), and axial plumes deeper (1706–1888 m) than the eruption (purple triangles, y = 4.2 ×10−16 + 1.5 × 10−17x, r2 = 0.88). The green short‐dashedline shows the 3He/heat ratio (1.6 × 10−17) from a singlesample collected from the Nautilus vent discharge in2009. The blue square in brackets denotes a sample col-lected by the CTD close to the Nautilus vent that wasnot used in the regression calculation. Large shaded areasshow the range of 3He/heat ratios for other confirmedevent plumes (red) and vent fluids in magmatically quies-cent sites (green) (see Table 2 for ratios).

Figure 10. DMn/heat ratios (nmol/J) from the 2008NELSC eruption site for samples in event plumes (red dots,y = 21 + 0.040x, r2 = 0.57), over the lava flow at depths of1500–1650 m (blue squares, y = −170 + 0.64x, r2 = 0.92),and axial plumes deeper (1706–1888 m) than the eruption(purple triangles, y = 9.0 + 0.41x, r2 = 0.96). NoDMn/heatratio is available from the Nautilus vent. The bracketed reddot is a sample of uncertain DMn quality not used in theregression calculation. Large shaded areas show the rangeof DMn/heat ratios for other confirmed event plumes (red)and vent fluids inmagmatically quiescent sites (green) (seeTable 2 for ratios).


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from ridge sources deeper than the eruption or on theadjoining NELSC segments (1706–1888 m) were10 times higher, comparable to chronic plumes justafter eruptions on othermid‐ocean ridges (Figure 11).Ratios from the Nautilus andMaka vents (sampled in2009) are even higher, at 1.6 to 5 × 10−17 mol/J.

[26] Unlike 3He/heat, DMn/heat ratios in the 2008event plumes were within the range of past eruptionplumes (Figure 10 and Table 2). Post‐eruption andnon‐eruption plumes had similar DMn/heat ratios,both falling in the middle of the broad range ofDMn/heat ratios from high‐temperature vents atmagmatically quiescent sites (Figure 10). Modestincreases in chronic plume DMn/heat have some-times accompanied past eruptions [Baker et al.,1995; Massoth et al., 1994], but even these increa-ses are small relative to geographic variability.

4.3.2. H2

[27] Concentrations of H2 have rarely been mea-sured in event plumes. Small concentrations of H2

in event plumes were first seen in EP96A and Bover the Gorda Ridge (up to 47 nM in EP96A)[Kelley et al., 1998]. These observations werenoteworthy because evidence suggests that H2 in

plumes results from water/rock reactions, particu-larly at high temperatures, wherein H2 is producedfrom the reduction of water by reduced iron com-pounds in the rock [Sansone et al., 1991; Seyfriedand Ding, 1995; Lilley et al., 2003; McCollomand Seewald, 2007]. In slow spreading environ-ments, H2 can be produced by the serpentinizationof mantle rocks [Charlou et al., 2002; Proskurowskiet al., 2006]. The 2008 event plumes were uniquein having consistently high (up to 9123 nM) H2

concentrations [Baumberger et al., 2009]. Unlike3He and DMn, H2 concentrations are highest inthe event plumes and lowest in the chronic plumes(Figure 8). All shallow plume samples had high butvariable H2/heat ratios (1.6 to 6.0 nmol/J). Thisvariability presumably arose from some combina-tion of differing extents of lava‐fluid interactionduring the formation of individual plume bursts,and progressive H2 loss by microbial oxidation. Thefew measurements available for microbial oxida-tion rates in hydrothermal plumes indicate H2 half‐lives ranging from a few hours to a few days[Kadko et al., 1990; Lilley et al., 1995;McLaughlin‐West et al., 1999]. The fact that samples fromadjacent event plumes in 2008 had roughly similarH2/heat ratios, and that these ratios did not decreasewith distance from the eruption site (i.e., time since

Table 2. Chemical Characteristics of Event Plumes and Vent Fluids

Source

3He/Heata

(10−17 mol/J)DMn/Heata

(nmol/J) Reference

Confirmed Event PlumesEP86 0.16 0.055 Baker et al. [1987, 1989]; Lupton et al. [1989]EP87A 0.16 0.040 Baker et al. [1989]EP93A 0.16 0.065 Lupton et al. [1995]; Massoth et al. [1995]EP93B 0.16 0.085 Lupton et al. [1995]; Massoth et al. [1995]EP93C 0.17 0.065 Lupton et al. [1995]; Massoth et al. [1995]EP96A 0.18 0.050 Kelley et al. [1998]; Massoth et al. [1998]EP96B 0.20 0.045 Kelley et al. [1998]; Massoth et al. [1998]EP08 0.041 0.040 This paper

Possible Event PlumesEP87B NDb 0.031 Nojiri et al. [1989]EP90 ND NDEP98 0.45 0.15 Lupton et al. [1999b]; Resing et al. [1999]EP01 ND 0.027 Edmonds et al. [2003]EP03 ND 0.080 Murton et al. [2006]

Vents at Magmatically Quiescent SitesGalapagos, 1977 0.52 0.6 Jenkins et al. [1978]21°N, East Pacific Rise, 1979 0.52 0.6 Lupton et al. [1980]; Welhan and Craig [1983]13°N, EPR, 1982, 1984 0.71–1.5 0.49–1.19 Merlivat et al. [1987]S Cleft, Juan de Fuca Ridge, 1984 0.53 3.65 Evans et al. [1988]; Kennedy [1988]; Massoth et al. [1994]Snake Pit, Mid‐Atlantic Ridge, 1986 0.6–1.3 0.22 Rudnicki and Elderfield [1992]TAG, MAR, 1993–1995 0.5–1.3 0.27 Rudnicki and Elderfield [1992]; Charlou et al. [1996]Lucky Strike, MAR, 1993 0.58 0.14 Jean‐Baptiste et al. [1998]

aPlume samples are corrected for stratification effects on D�.bND, not determined.


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discharge) (Figure 6), suggests that variability insource fluid H2 was more important than oxidationthrough time in creating the event plume variabilityin H2/heat ratios.

[28] We suggest that the 100‐fold difference inmaximum H2 concentrations (and in H2/heat ratio)between EP96A [Kelley et al., 1998] and EP08 isdominantly a function of eruption style, not H2

oxidation, because both event plumes were ofroughly similar ages. Based on the seismic record[Fox and Dziak, 1998], EP96A was most likely∼1 day old when discovered, and certainly no more

than 10 days old. This young age is consistent withan age estimate of <3 days based on the ratio ofdissolved to particulate Fe in the plume [Massothet al., 1998]. Clague et al. [2009] have now docu-mented that an almost unfailing characteristic ofsubmarine eruptions, sampled throughout the Pacificat depths from 1400 to 3800 m, is the productionof glassy pyroclastic fragments that are remnantsof bubbles of magmatic gas. Such fragments wereabundant on and around the Puipui flow on theNELSC [Clague et al., 2010], and imply extensiveopportunities for lava‐seawater interaction during

Figure 11. Time series of 3He/heat ratios (mol/J) in the chronic plumes (dots) and in fluids from sulfide chimneys(triangles) at various eruption sites. The “Quiescent” values include plume measurements made pre‐eruption at aneruption site, >1000 days post‐eruption, or at a vent field(s) near but not at an eruption. Note that only the 2008NELSC data show very low eruption‐associated 3He/heat ratios. Sources: Axial volcano, Lupton et al. [1999b]and Resing et al. [2004]; Cleft, Baker and Lupton [1990]; EPR 9.9°N, Lupton et al. [1993]; Gorda Ridge, Kelleyet al. [1998]; CoAxial, Lupton et al. [1995, also unpublished data, 2011]; vent fluids, Lilley et al. [2003]; Lau,this paper. The “Quiescent” value (open square) is the mean and standard deviation of vent fluid 3He/heat in Table 2.


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the eruption. Of the dozens of eruption sites thatClague et al. [2009] sampled in the Pacific, the 1996Gorda Ridge site alone was without pyroclasticdebris.

[29] All other NELSC plume samples, whetherfrom just above the lava field or from venting muchdeeper than the eruption, had uniformly low H2/heat ratios (<0.41 nmol/J, Figure 8), as also foundin chronic venting from the Gorda Ridge eruption[Kelley et al., 1998]. Fluids supplying these plumeshad much less, if any, contact with molten lava.

[30] We can assess the scale of H2 water‐rockinteraction by estimating the volume of rock neededto supply the observed H2. In the 2008 event plumes,H2 andD� are linearly correlated (r2 = 0.68), so fromthe total hydrothermal heat inventory in the plumes(1.2 × 1015 J) we estimate the total H2 at 4.4 × 10

6M.To produce this quantity of H2 an equivalent molaramount of water is needed, which corresponds to7.9 × 104 kg, or 79 m3, of H2O. We assume that83% of Fe in the basalts is FeO, which is themedian Fe2+/(FeTOT) value of six Lau Basin basaltshaving similar major element compositions and aFe2+/(FeTOT) range of 70 to 90% [Nilsson and Peach,1993]. A similar molar quantity of FeO, or 3.2 ×105 kg, must react with the H2O. Analysis of thePuipui basalts shows a FeO abundance of 9.1%,thus requiring 3.5 × 106 kg, or ∼1200 m3, of basalt.The area of the flow, based on ∼14 h of ROVmapping, is 0.6–0.9 km2. These are likely minimumvalues, as the off‐axis extent of the flow wasinvestigated in only one region. Estimates of theflow volume from JASON observations [Rubin et al.,2009] and bathymetry differencing between 2006and 2009 (W. W. Chadwick, Jr., personal commu-nication, 2011) yield a flow volume of ∼1.5 to 5.8 ×106 m3. Only 0.08–0.02% of the lava had to reactwith seawater to produce the H2 inventory in theevent plumes (assuming the entire lava volume wasemplaced at one time). Because this small volumeis equivalent to an average thickness of ∼1–2 mmover the flow area, it is consistent with H2 formingprimarily during the fraction of a second requiredfor the surface (i.e., the outermost 1–2 mm) ofmolten lava to attain near‐ambient temperature[Gregg and Fornari, 1998].

[31] Another indication that the event plume sourcefluids experienced lava‐seawater interaction was thediscovery of abundant glassy shards in the plumes(Figure 6). Such fragments are commonly foundin association with submarine eruptions [Maicherand White, 2001; Eissen et al., 2003; Clague et al.,2009]. The largest shards sampled in 2008 hadStokesian fall velocities on the order of 100 m/d andwere found as high as 600 m above the depth of the

eruption zone. The presence of such particles in thin(∼60 m) plumes is consistent with the presumptionthat these plumes were no more than a few days oldwhen observed.

5. The Role of Lava in Event PlumeFormation

[32] The 2008 NELSC eruption extends the corre-lation between event plume heat anomaly and lavaheat content across a factor of ∼100 in both vari-ables (Figure 12). The event plume heat as a per-centage of total available lava heat varies widely,from 4 to 160%. This range could vary becausethere is no causal relationship between the two,because of the uncertainties in accurately deter-mining the volumes of event plumes and lava flows,or because the percentage of lava heat transferred toevent plumes has varied among sites. If event plumeheat derives from lava cooling, we would expectboth a correlation between the two heat inventoriesand the lava heat inventory to always exceed theevent plume inventory. Uncertainties arise in thesecalculations, of course, because underestimation ofevent plume heat is always possible (e.g., unde-tected event plumes) and the lava heat determina-tion depends on an accurate knowledge of thelava volume. Nevertheless, the fact that the 1993CoAxial event plumes held >100% of the calcu-lated lava heat (assuming the lava volume has notbeen underestimated by more than ∼50%) is achallenge for the lava‐cooling hypothesis.

[33] If event plumes do form solely by the coolingof erupted lava [Palmer and Ernst, 1998, 2000],most of the available lava heat (Figure 12) must betransferred fast enough to satisfy the short timeconstraints of event plume formation. Models ofevent plume creation [Lavelle, 1995] and measure-ments of the precipitation rate of dissolved Fe inevent plumes [Massoth et al., 1995, 1998] provideconvincing evidence that event plume formation timeis hours rather than many days. Our observations offast‐settling basalt shards, high H2 concentrations,high dE/dt anomalies, and rapid advection of theobserved plumes out of the eruption area also requirea short formation interval.

[34] Although there are no direct measurements ofthe in situ cooling rate of solidifying submarinelavas, theoretical calculations and field observa-tions provide reliable estimates of typical coolingrates. Gregg and Fornari [1998] calculated thegrowth rate of the solid crust (defined as the depth


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of the glass transition temperature (732°C)) of asubmarine lava flow and found the rate to be pro-portional to the square root of time (also derived byGriffiths and Fink [1992]). Ideally cooling lavacrust reaches a 0.015 m thickness in ∼40 s, 0.1 m in∼30 min, and 1 m in ∼48 h. Cooling to 732°Creleases about half the original lava heat. The totalheat lost from a thickness of cooling crust is some-what greater, however, because the solidified crustcontinues to cool as the 732°C isotherm progressesinward.

[35] The validity of this theoretical calculation canbe tested against a real lava flow using data from ahigh‐precision pressure sensor serendipitouslyencased within the 1998 sheet flow eruption at

Axial volcano [Fox et al., 2001]. Pressure changesshowed that initial inflation (during lava extrusion)lasted 72 min, followed by 81 min of steady defla-tion and 44min of a secondminor inflation/deflationevent. Thus for at least 3 h and 17 min after theeruption began, the thin sheet flow remained fluidenough to inflate 3.5 m and deflate 2.5 m. Chadwick[2003] modeled the drainback process by measur-ing the thickness of drainback cavity lids (∼0.1 m)and crustal shelves (∼0.015 m) on lava pillarswithin the cavities. These thicknesses and theirformation time (∼1 h and 24 s, respectively), cal-culated from the pressure sensor record, agree withthe theoretical growth rate of lava crust calculatedbyGregg and Fornari [1998]. (Themeasured growth

Figure 12. Scatterplot and linear regression between event plume heat and lava heat for the four eruptions wheresuch data are available (EP86, 93, 96, and 08; see Table 1). For the high estimate of lava volume at each eruption(red line and circles) the plume heat represents 70% of the lava heat (y = 1.03 × 1013 + 0.70x, r2 = 0.83). For the lowestimate (blue line and dots) the plume heat represents 95% of the lava heat (y = 1.97 × 1015 + 0.95x, r2 = 0.93). Notethat the 1993 CoAxial event plume heat (0.7 × 1017 J) exceeds the calculated available lava heat at that site.


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time of the cavity lid (1 h) is somewhat longer thanthe theoretical value (0.5 h) because of additionalheat supplied by the flowing lava and uncertaintiesin applying theory to actual rock.) Moreover, thefact that most of the emplaced lava volumeremained fluid, and thus hotter than ∼1000°C, for atleast 3 h after the eruption began restricts the lavaheat loss to only ∼10–25% of its total content(depending on the extent of crystallization in thelava).

[36] An additional proposed cooling mechanism islava fountaining during volatile‐rich eruptions ofthe Strombolian or Hawaiian style [e.g., Head andWilson, 2003]. Despite the near‐ubiquitous asso-ciation of quenched pyroclastic debris and erup-tions, however, pyroclastic fragments apparentlyaccount for only a negligible mass fraction of theerupted lava even when they appear “abundant” onthe seafloor. The mass fraction of pyroclasts hasbeen quantified at only one site, “NESCA” in theGorda Ridge Escanaba Trough, where it totaled<0.02% of the total lava mass [Clague et al., 2009].Thus their rapid quenching can transfer only anegligible fraction of the lava heat to seawater.

[37] These results indicate that even for thin sheetflows (as at the NELSC) near‐complete cooling in<∼1 day is improbable. The case for rapid heattransfer at other sites is even more problematical.Eruptions at Cleft, CoAxial, and Gorda Ridge allproduced thick (25–75 m) constructions of pillowmounds that likely required several days to severalweeks to complete [Chadwick and Embley, 1994;Gregg and Fink, 1995; Chadwick et al., 1998;Rubin et al., 1998]. At the Gorda Ridge site, thepillow mound is unusually thick (up to 75 m) andsteep in some places, thin (∼10 m) elsewhere, andevidently emplaced slowly over several weeks[Chadwick et al., 1998; Rubin et al., 1998]. Thiscomplicated mound topography bears little resem-blance to the ideal layer‐by‐layer, emplacement‐then‐cooling construction required by the model ofPalmer and Ernst [1998, 2000].Clague et al. [2009]note that the absence of pyroclastic debris at theGorda Ridge site (about the same depth, ∼3200 m,as the NESCA site) is another indication of a slowextrusion rate eruption. (The absence of pyroclasticdebris here also weakens the proposal of Clagueet al. [2009] that gas bubbles are the primarysource of magmatic gases (e.g., 3He) found in allevent plumes.)

[38] Conductive heat flow observations provide alonger view of pillow mound cooling rates. Mea-surements at the summit of the thickest portion of

the Gorda Ridge mound eight months after theeruption found heat flow (up to 6950 mW/m2)at least 100 times higher than on “old” crust justadjacent to other lava flows [Johnson and Hutnak,1997]. Combining the Gorda Ridge results withsimilar measurements at the CoAxial eruptionmound, Johnson and Hutnak [1997] suggest thatthe cooling half‐life of those flows was ∼2 yr, fartoo long to power event plumes.

[39] The Gorda Ridge flow may be an extremeexample of a slowly cooling lava eruption, but thissite nevertheless produced two event plumes. Thefirst (EP96A) was discovered 11 days after regionalseismicity began and only 1 day after the onset of a10‐day long seismic event centered at the locationof the lava mound and EP96A [Fox and Dziak,1998]. Thus, even given a convincing relationshipbetween erupted lava heat and event plume heat,any theory that postulates lava cooling as the directsource of event plume buoyancy flux must accountfor rapid heat transfer from flows ranging fromthin, rapidly erupted sheets to thick, slowly extrudedpillow mounds.

6. Summary

[40] The discovery of an eruption event on theNELSC in 2008 provided the first new clues in thepuzzle of event plume formation in more than adecade. We now know that event plumes with achemically unique and uniform composition occurin volumes spanning three orders of magnitude,that event plume source fluids interact enough withmolten lava to generate very high concentrations ofH2, and that the anomalous heat content in eventplumes is roughly equivalent (∼70–95% on aver-age) to the heat content of the erupted lavas inexamples spanning two orders of magnitude. Theseresults seem to point to the generation of eventplumes by a rapid transfer of heat and chemicalsfrom lava to seawater, but there are serious physicalconstraints to this hypothesis. Theoretical calcula-tions and field observations of cooling lava flows,both thin sheet flows and thick pillow mounds, allconclude that erupted lava cools far more slowlythan required to create the intense buoyancy fluxneeded to lift voluminous plumes up to a kilometerabove the seafloor [e.g., Lavelle, 1995]. If lavacooling does generate event plumes, then we mustbe quite ignorant about how lava actually cools onthe deep seafloor.

[41] Heat source is one key to understanding eventplumes; the other is fluid chemistry (see Lupton


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et al. [1999a, 2000] and Palmer and Ernst [1998,2000] for a thorough discussion). Although thispaper does not focus on the origin of event plumefluid chemistry, the 2008 event did provide somenew chemical constraints. First, unlike every othersampled event, the near‐bottom chronic plumesassociated with the erupted lava had 3He/heat ratiosmuch lower than nearby plumes not affected by theeruption. Those ratios were also much lower thanhistorical post‐eruption plumes. It is possible thatthe low NELSC 3He/heat ratios represented anevolving transition from event plume dischargeto typical post‐eruption discharge, but if so, thattransitionmust have beenmuch slower than observedfollowing other events. Second, the extraordinarilyhigh H2 concentrations in the event plumes appar-ently demand significant lava‐fluid interactionregardless of the source of the fluids. And third, thecombination of event plume D� anomalies com-parable to prior event plumes and a relatively highd�/dS gradient in the Lau Basin virtually eliminatesthe possibility that the source of event plumes arelow‐salinity fluids such as sampled immediatelypost‐eruption at other locations [Von Damm et al.,1995; Lilley et al., 2003].

Notation

DNTU Relative measure of light backscatteringabove ambient (dimensionless).

E Relative measure of oxidation‐reductionpotential (mV).

�, �v Potential temperature of seawater or ventfluid (°C).

D� Hydrothermal temperature anomaly (°C).s� Potential density ((kg/m3)‐1000).

S, Sv Salinity of seawater or vent fluid (psu).Q, Qv Volume flux of plume or vent discharge

(cm3/s4).A Coefficient of thermal expansion (1/°C).b Coefficient of haline contraction (1/psu).

Acknowledgments

[42] This research was sponsored by the NOAA VENTS Pro-gram, the NOAA Office of Ocean Exploration and Research,and NSF Marine Geology and Geophysics. D. A. Clague,G. J. Massoth, and J. W. Lavelle provided valuable discussionand comments for improving the manuscript. ReviewersT. Gregg, J.‐L. Charlou, and an Associate Editor also offereddetailed and constructive suggestions. This paper is PMEL con-tribution 3682.

References

American Public Health Association (1985), Standard Methodsfor the Examination of Water and Wastewater, 16th ed.,1268 pp., Am. Public Health Assoc., Washington, D. C.

Baker, E. T. (1998), Patterns of event and chronic hydrothermalventing following a magmatic intrusion: New perspectivesfrom the 1996 Gorda Ridge eruption, Deep Sea Res., Part II,45, 2599–2618, doi:10.1016/S0967-0645(98)00085-X.

Baker, E. T., and J. L. Lupton (1990), Changes in sub-marine hydrothermal helium‐3/heat ratios as an indicatorof magmatic/tectonic activity, Nature, 346, 556–558,doi:10.1038/346556a0.

Baker, E. T., G. J.Massoth, andR. A. Feely (1987), Cataclysmichydrothermal venting on the Juan de Fuca Ridge,Nature, 329,149–151, doi:10.1038/329149a0.

Baker, E. T., J. W. Lavelle, R. A. Feely, G. J. Massoth, S. L.Walker, and J. E. Lupton (1989), Episodic venting of hydro-thermal fluids from the Juan de Fuca Ridge, J. Geophys.Res., 94, 9237–9250, doi:10.1029/JB094iB07p09237.

Baker, E. T., G. J. Massoth, R. A. Feely, R. W. Embley, R. E.Thomson, and B. J. Burd (1995), Hydrothermal event plumesfrom the CoAxial seafloor eruption site, Juan de Fuca Ridge,Geophys. Res. Lett., 22, 147–150, doi:10.1029/94GL02403.

Baker, E. T., C. G. Fox, and J. P. Cowen (1999), In situ obser-vations of the onset of hydrothermal discharge during the1998 submarine eruption of Axial Volcano, Juan de FucaRidge, Geophys. Res. Lett., 26, 3445–3448, doi:10.1029/1999GL002331.

Baker, E. T., F. Martinez, J. A. Resing, S. L. Walker, N. J.Buck, and M. H. Edwards (2010), Hydrothermal coolingalong the Eastern Lau Spreading Center: No evidence fordischarge beyond the neovolcanic zone, Geochem. Geophys.Geosyst., 11, Q08004, doi:10.1029/2010GC003106.

Baumberger, T., M. D. Lilley, J. E. Lupton, J. A. Resing, E. T.Baker, S. L. Walker, K. H. Rubin, and G. L. Frueh‐Green(2009), Hydrogen and methane dissolved in plumes of theNortheast Lau Basin eruption sites, Eos Trans. AGU, 90(52),Fall Meet. Suppl., Abstract V43I‐04.

Butterfield, D. A., et al. (1997), Seafloor eruptions and evolu-tion of hydrothermal fluid chemistry, Philos. Trans. R. Soc.London A, 355, 369–386, doi:10.1098/rsta.1997.0013.

Cann, J. R., and M. R. Strens (1989), Modeling periodic mega-plume emission by black smoker systems, J. Geophys. Res.,94, 12,227–12,237, doi:10.1029/JB094iB09p12227.

Chadwick, W. W., Jr. (2003), Quantitative constraints on thegrowth of submarine lava pillars from amonitoring instrumentthat was caught in a lava flow, J. Geophys. Res., 108(B11),2534, doi:10.1029/2003JB002422.

Chadwick, W. W., Jr., and R. W. Embley (1994), Lava flowsfrom a mid‐1980s submarine eruption on the Cleft segment,Juan de Fuca Ridge, J. Geophys. Res., 99, 4761–4776,doi:10.1029/93JB02041.

Chadwick, W. W., Jr., R. W. Embley, and C. G. Fox (1991),Evidence for volcanic eruption on the southern Juan de FucaRidge between 1981 and 1987, Nature, 313, 212–214.

Chadwick, W. W., Jr., R. W. Embley, and C. G. Fox (1995),SeaBeam depth changes associated with recent lava flows,CoAxial segment, Juan de Fuca Ridge: Evidence for multi-ple eruptions between 1981–1993, Geophys. Res. Lett., 22,167–170, doi:10.1029/94GL01895.

Chadwick, W. W., Jr., R. W. Embley, and T. M. Shank (1998),The 1996 Gorda Ridge eruption: Geologic mapping, sidescansonar, and SeaBeam comparison results,Deep SeaRes., Part II,45, 2547–2569, doi:10.1016/S0967-0645(98)00083-6.


18 of 21

Charlou, J. L., J. P. Donval, P. Jean‐Baptiste, A. Dapoigny,and P. A. Rona (1996), Gases and helium isotopes in hightemperature solutions sampled before and after ODP Leg158 drilling at TAG Hydrothermal Field (26°N‐MAR),Geophys. Res. Lett., 23, 3491–3494, doi:10.1029/96GL02141.

Charlou, J. L., J. P. Donval, Y. Fouquet, P. Jean‐Baptiste, andN. Holm (2002), Geochemistry of high H2 and CH4 ventfluids issuing from ultramafic rocks at the Rainbow hydro-thermal field (36°14′N, MAR), Chem. Geol., 191, 345–359,doi:10.1016/S0009-2541(02)00134-1.

Clague, D. A., J. B. Paduan, and A. S. Davis (2009), Wide-spread strombolian eruptions of mid‐ocean ridge basalt,J. Volcanol. Geotherm. Res., 180, 171–188, doi:10.1016/j.jvolgeores.2008.08.007.

Clague, D. A., D. W. Caress, K. H. Rubin, and J. B. Paduan(2010), The 2008 Puipui eruption and morphology of theNortheast Lau Spreading Center between Maka and Tafu,Abstract T11E‐03 presented at 2010 Fall Meeting, AGU,San Francisco, Calif., 13–17 Dec.

Dziak, R. P., D. R. Bohnenstiehl, J. P. Cowen, E. T. Baker, K. H.Rubin, J. H. Haxel, and M. J. Fowler (2007), Rapid dikeemplacement leads to dramatic hydrothermal plume releaseduring seafloor spreading events, Geology, 35, 579–582,doi:10.1130/G23476A.1.

Edmonds, H. N., et al. (2003), Discovery of abundant hydro-thermal venting on the ultraslow‐spreading Gakkel Ridge,ArcticOcean,Nature, 421, 252–256, doi:10.1038/nature01351.

Eissen, J.‐P., Y. Fouquet, D. Hardy, and H. Ondreas (2003),Recent MORB volcaniclastic explosive deposits formedbetween 500 and 1750 m.b.s.l. on the axis of the Mid‐AtlanticRidge, south of the Azores, in Explosive Subaqueous Volca-nism, Geophys. Monogr. Ser., vol. 140, edited by J. D. L.White, J. L. Smellie, and D. A. Clague, pp. 143–166, AGU,Washington, D. C.

Embley, R. W., W. W. Chadwick Jr., D. Clague, and D. Stakes(1999), 1998 eruption of Axial Volcano: Multibeam anoma-lies and seafloor observations, Geophys. Res. Lett., 26,3425–3428, doi:10.1029/1999GL002328.

Evans, W., L. White, and J. Rapp (1988), Geochemistry ofsome gases in hydrothermal fluids from the southern Juande Fuca Ridge, J. Geophys. Res., 93(B12), 15,305–15,313,doi:10.1029/JB093iB12p15305.

Fox, C. G., and R. P. Dziak (1998), Hydroacoustic detection ofvolcanic activity on the Gorda Ridge, February–March 1996,Deep Sea Res., Part II, 45, 2513–2530, doi:10.1016/S0967-0645(98)00081-2.

Fox, C. G., W. W. Chadwick Jr., and R. W. Embley (2001),Direct observation of a submarine volcanic eruption from asea‐floor instrument caught in a lava flow, Nature, 412,727–729, doi:10.1038/35089066.

Gamo, T., H. Sakai, J. Ishibashi, E. Nakayama, K. Isshiki,H.Matsuura, K. Shitashima, K. Takeuchi, and S. Ohta (1993),Hydrothermal plumes in the eastern Manus Basin, BismarkSea: CH4, Mn, Al, and pH anomalies, Deep Sea Res., 40,2335–2349, doi:10.1016/0967-0637(93)90108-F.

German, C. R., E. T. Baker, D. P. Connelly, J. E. Lupton,J. Resing, R. D. Prien, S. L. Walker, H. N. Edmonds, andC. H. Langmuir (2006), Hydrothermal exploration of theFonualei Rift and Spreading Center and the Northeast LauSpreading Center, Geochem. Geophys. Geosyst., 7, Q11022,doi:10.1029/2006GC001324.

Gregg, T. K. P., and J. H. Fink (1995), Quantification ofsubmarine lava‐flow morphology through analog experi-ments, Geology, 23, 73–76, doi:10.1130/0091-7613(1995)023<0073:QOSLFM>2.3.CO;2.

Gregg, T. K. P., and D. J. Fornari (1998), Long submarine lavaflows: Observations and results from numerical modeling,J. Geophys. Res., 103, 27,517–27,531, doi:10.1029/98JB02465.

Griffiths, R. W., and J. H. Fink (1992), Solidification and mor-phology of submarine lavas: A dependence on extrusion rate,J. Geophys. Res., 97, 19,729–19,737, doi:10.1029/92JB01594.

Hawkins, J. W. (1995), Evolution of the Lau Basin—Insightsfrom ODP Leg 135, in Active Margins and Marginal Basinsof the Western Pacific, Geophys. Monogr. Ser., vol. 188,edited by B. Taylor and J. P. Natland, pp. 125–173, AGU,Washington, D. C.

Head, J. W., III, and L. Wilson (2003), Deep submarine pyro-clastic eruptions: Theory and predicted landforms anddeposits, J. Volcanol. Geotherm. Res., 121, 155–193,doi:10.1016/S0377-0273(02)00425-0.

Jean‐Baptiste, P., H. Bougault, A. Vangriesheim, J. L. Charlou,J. Radford‐Knoery, Y. Fouquet, D. Needham, and C. German(1998), Mantle 3He in hydrothermal vents and plume of theLucky Strike site (MAR 37°17′N) and associated geothermalheat flux, Earth Planet. Sci. Lett., 157, 69–77, doi:10.1016/S0012-821X(98)00022-3.

Jenkins, W. J., J. M. Edmond, and J. B. Corliss (1978), Excess3He and 4He in Galapagos submarine hydrothermal waters,Nature, 272, 156–158, doi:10.1038/272156a0.

Johnson, P., and M. Hutnak (1997), Conductive heat loss inrecent eruptions at mid‐ocean ridges, Geophys. Res. Lett.,24, 3089–3092, doi:10.1029/97GL02998.

Kadko, D., N. Rosenberg, J. Lupton, R. Collier, and M. D.Lilley (1990), Chemical reaction rates and entrainment withinthe Endeavour Ridge hydrothermal plume, Earth Planet. Sci.Lett., 99, 315–335, doi:10.1016/0012-821X(90)90137-M.

Kelley, D. S., M. D. Lilley, J. E. Lupton, and E. J. Olson(1998), Enriched H2, CH4, and

3He concentrations in hydro-thermal plumes associated with the 1996 Gorda Ridge eruptiveevent, Deep Sea Res., Part II, 45, 2665–2682, doi:10.1016/S0967-0645(98)00088-5.

Kennedy, B. M. (1988), Nobel gases in vent water from theJuan de Fuca Ridge, Geochim. Cosmochim. Acta, 52,1929–1935, doi:10.1016/0016-7037(88)90016-6.

Kim, J., S. K. Son, J. W. Son, K. H. Kim, W. J. Shim, C. H.Kim, and K. Y. Lee (2009), Venting sites along the Fonualeiand Northeast Lau spreading centers and evidence of hydro-thermal activity at an off‐axis caldera in the northeastern LauBasin, Geochem. J., 43, 1–13, doi:10.2343/geochemj.0.0164.

Lavelle, J. W. (1995), The initial rise of a hydrothermal plumefrom a line segment source—Results from a three‐dimen-sional numerical model, Geophys. Res. Lett., 22, 159–162,doi:10.1029/94GL01463.

Lavelle, J. W., E. T. Baker, and G. J. Massoth (1998), On thecalculation of total heat, salt, and tracer fluxes from oceanhydrothermal events, Deep Sea Res., Part II, 45, 2619–2636,doi:10.1016/S0967-0645(98)00086-1.

Lilley, M. D., R. A. Feely, and J. H. Trefry (1995), Chemicaland biochemical transformations in hydrothermal plumes, inSeafloor Hydrothermal Systems: Physical, Chemical, Biolog-ical, and Geological Interactions, Geophys. Monogr. Ser.,vol. 91, edited by S. Humphris et al., pp. 369–391, AGU,Washington D. C.

Lilley, M. D., D. A. Butterfield, J. E. Lupton, and E. J. Olson(2003), Magmatic events can produce rapid changes inhydrothermal vent chemistry, Nature, 422, 878–881,doi:10.1038/nature01569.


19 of 21

Lowell, R. P., and L. N. Germanovich (1995), Dike injectionand the formation of megaplumes at ocean ridges, Science,267, 1804–1807, doi:10.1126/science.267.5205.1804.

Lupton, J. E. (1990),Water column hydrothermal plumes on theJuan de Fuca Ridge, J. Geophys. Res., 95, 12,829–12,842,doi:10.1029/JB095iB08p12829.

Lupton, J. E., G. P. Klinkhammer, W. R. Normark, R. Haymon,K. C.Macdonald, R. F.Weiss, and H. Craig (1980), Helium‐3and manganese at the 21°N East Pacific Rise hydrothermalsite, Earth Planet. Sci. Lett., 50, 115–127, doi:10.1016/0012-821X(80)90123-5.

Lupton, J. E., E. T. Baker, and G. J. Massoth (1989), Variable3He/heat ratios in submarine hydrothermal systems: Evi-dence from two plumes over the Juan de Fuca Ridge, Nature,337, 161–164, doi:10.1038/337161a0.

Lupton, J. E., E. T. Baker, M. J. Mottl, F. J. Sansone, C. G.Wheat, J. A. Resing, G. J. Massoth, C. I. Measures, andR. A. Feely (1993), Chemical and physical diversity of hydro-thermal plumes along the East Pacific Rise, 8°45′N to11°50′N, Geophys. Res. Lett., 20, 2913–2916, doi:10.1029/93GL00906.

Lupton, J. E., E. T. Baker, G. J. Massoth, R. E. Thomson, B. J.Burd, D. A. Butterfield, R. W. Embley, and G. A. Cannon(1995), Variations in water‐column 3He/heat ratios asso-ciated with the 1993 CoAxial event, Juan de Fuca Ridge,Geophys. Res. Lett., 22, 155–158, doi:10.1029/94GL02797.

Lupton, J. E., E. T. Baker, and G. J. Massoth (1999a), Helium,heat, and the generation of hydrothermal event plumes atmid‐ocean ridges, Earth Planet. Sci. Lett., 171, 343–350,doi:10.1016/S0012-821X(99)00149-1.

Lupton, J. E., E. Baker, R. Embley, R. Greene, and L. Evans(1999b), Anomalous helium and heat signatures associatedwith the 1998 Axial Volcano event, Juan de Fuca Ridge,Geophys. Res. Lett. , 26 , 3449–3452, doi:10.1029/1999GL002330.

Lupton, J. E., E. T. Baker, and G. J. Massoth (2000), Reply toComment by M. R. Palmer and G. G. J. Ernst on “Helium,heat, and the generation of hydrothermal event plumesat mid‐ocean ridges” by J. E. Lupton, E. T. Baker, andG. J. Massoth, Earth Planet. Sci. Lett., 180, 219–222,doi:10.1016/S0012-821X(00)00151-5.

Maicher, D., and J. D. L. White (2001), The formation ofdeep‐sea Limu o Pele, Bull. Volcanol., 63, 482–496.

Massoth, G. J., E. T. Baker, J. E. Lupton, R. A. Feely, D. A.Butterfield, K. Von Damm, K. K. Roe, and G. T. Lebon(1994), Temporal and spatial variability of hydrothermalmanganese and iron at Cleft segment, Juan de Fuca Ridge,J. Geophys. Res., 99, 4905–4923, doi:10.1029/93JB02799.

Massoth, G. J., E. T. Baker, R. A. Feely, D. A. Butterfield, R. E.Embley, J. E. Lupton, R. E. Thomson, and G. A. Cannon(1995), Observations of manganese and iron at the CoAxialseafloor eruption site, Juan de Fuca Ridge, Geophys. Res.Lett., 22, 151–154, doi:10.1029/94GL02662.

Massoth, G. J., E. T. Baker, R. A. Feely, J. E. Lupton, R. W.Collier, J. F. Gendron, K. K. Roe, S. M. Maenner, and J. A.Resing (1998), Manganese and iron in hydrothermal plumesresulting from the 1996 Gorda Ridge event, Deep Sea Res.,Part II, 45, 2683–2712, doi:10.1016/S0967-0645(98)00089-7.

McCollom, T. M., and J. S. Seewald (2007), Abiotic synthesisof organic compounds in deep‐sea hydrothermal environ-ments, Chem. Rev., 107, 382–401, doi:10.1021/cr0503660.

McDougall, T. J. (1990), Bulk properties of “hot smoker”plumes, Earth Planet. Sci. Lett., 99, 185–194, doi:10.1016/0012-821X(90)90081-8.

McDuff, R. E. (1995), Physical dynamics of deep‐sea hydro-thermal plumes, in Seafloor Hydrothermal Systems: Physi-cal, Chemical, Biological, and Geological Interactions,Geophys. Monogr. Ser., vol. 91, edited by S. Humphris et al.,pp. 357–368, AGU, Washington D. C.

McLaughlin‐West, E. A., E. J. Olson,M. D. Lilley, J. A. Resing,J. E. Lupton, E. T. Baker, and J. P. Cowen (1999), Variationsin hydrothermal methane and hydrogen concentrations fol-lowing the 1998 eruption at Axial Volcano, Geophys. Res.Lett., 26, 3453–3456, doi:10.1029/1999GL002336.

Merlivat, L., F. Pineau, and M. Javoy (1987), Hydrothermalvent waters at 13°N on the East Pacific Rise: Isotopic com-position and gas concentration, Earth Planet. Sci. Lett., 84,100–108, doi:10.1016/0012-821X(87)90180-4.

Murton, B. J., E. T. Baker, C. M. Sands, and C. R. German(2006), Detection of an unusually large hydrothermal eventplume above the slow‐spreading Carlsberg Ridge: NW IndianOcean, Geophys. Res. Lett., 33, L10608, doi:10.1029/2006GL026048.

Nakamura, K., S. Veirs, C. P. Sarason, R. E. McDuff, F. Stahr,D. R. Yoerger, and A. M. Bradley (2000), Electrochemicalsignals in rising buoyant plumes and tidally oscillating plumesat the Main Endeavour vent field, Juan de Fuca Ridge, EosTrans. AGU, 81(48), Fall Meet. Suppl., Abstract OS52I‐05.

Nilsson, K., and C. L. Peach (1993), Sulfur speciation, oxidation‐state, and sulfur concentration in backarc magmas, Geochim.Cosmochim. Acta, 57, 3807–3813, doi:10.1016/0016-7037(93)90158-S.

Nojiri, Y., J. Ishibashi, T. Kawai, and H. Sakai (1989), Hydro-thermal plumes along the North Fiji Basin spreading axis,Nature, 342, 667–670, doi:10.1038/342667a0.

Palmer, M. R., and G. G. J. Ernst (1998), Generation of hydro-thermal megaplumes by cooling of pillow basalts at mid‐ocean ridges, Nature, 393, 643–647, doi:10.1038/31397.

Palmer, M. R., and G. G. J. Ernst (2000), Comment on“Helium, heat, and the generation of hydrothermal eventplumes at mid‐ocean ridges”, by J. E. Lupton, E. T. Baker,and G. J. Massoth, Earth Planet. Sci. Lett., 180, 215–218,doi:10.1016/S0012-821X(00)00150-3.

Pelletier, B., S. Calmant, and R. Pillet (1998), Current tectonicsof the Tonga‐New Hebrides region, Earth Planet. Sci. Lett.,164, 263–276, doi:10.1016/S0012-821X(98)00212-X.

Proskurowski, G., M. D. Lilley, D. S. Kelley, and E. J. Olson(2006), Low temperature volatile production at the Lost CityHydrothermal Field, evidence from a hydrogen stable iso-tope geothermometer, Chem. Geol. , 229 , 331–343,doi:10.1016/j.chemgeo.2005.11.005.

Resing, J. A., and R. W. Embley (2009), Active eruptions inthe NE Lau Basin, Eos Trans. AGU, 90(52), Fall Meet.Suppl., Abstract V431‐01.

Resing, J. A., R. A. Feely, G. J. Massoth, and E. T. Baker(1999), The water column chemical signature after the1998 eruption of Axial Volcano, Geophys. Res. Lett., 26,3645–3648, doi:10.1029/1999GL002350.

Resing, J. A., J. E. Lupton, R. A. Feely, and M. D. Lilley(2004), CO2 and

3He in hydrothermal plumes: Implicationsfor mid‐ocean ridge CO2 flux, Earth Planet. Sci. Lett.,226, 449–464, doi:10.1016/j.epsl.2004.07.028.

Resing, J. A., G. Lebon, E. T. Baker, J. E. Lupton, R.W. Embley,G. J. Massoth, W. W. Chadwick Jr., and C. E. J. de Ronde(2007), Venting of acid‐sulfate fluids in a high‐sulfidation set-ting at NW Rota‐1 submarine volcano on the Mariana Arc,Econ. Geol., 102, 1047–1061, doi:10.2113/gsecongeo.102.6.1047.


20 of 21

Rubin, K. H., M. C. Smith, M. R. Perfit, D. M. Christie, andL. F. Sacks (1998), Geochronology and geochemistry oflavas from the 1996 North Gorda Ridge eruption, DeepSea Res., Part II, 45, 2571–2597, doi:10.1016/S0967-0645(98)00084-8.

Rubin, K. H., R. W. Embley, D. A. Clague, J. A. Resing, P. J.Michael, N. S. Keller, and E. T. Baker (2009), Lavas fromactive boninite and very recent basalt eruptions at two sub-marine NE Lau Basin sites, Eos Trans. AGU, 90(52), FallMeet. Suppl., Abstract V431‐05.

Rudnicki, M. D., and H. Elderfield (1992), Helium, radon, andmanganese at the TAG and Snakepit hydrothermal ventfields, 26° and 23°N, Mid‐Atlantic Ridge, Earth Planet.Sci. Lett., 113, 307–321, doi:10.1016/0012-821X(92)90136-J.

Sansone, F. J., J. A. Resing, G. W. Tribble, P. N. Sedwick,K. M. Kelly, and K. Hon (1991), Lava‐seawater interac-tions at shallow‐water submarine lava flows, Geophys.Res. Lett., 18, 1731–1734, doi:10.1029/91GL01279.

Seyfried, W. E., Jr., and K. Ding (1995), Phase equilibria insubseafloor hydrothermal systems: A review of the role ofredox, temperature, pH and dissolved Cl on the chemistryof hot spring fluids at Mid‐Ocean Ridges, in SeafloorHydrothermal Systems: Physical, Chemical, Biological, andGeological Interactions, Geophys. Monogr. Ser., vol. 91,edited by S. Humphris et al., pp. 248–272, AGU, Washington,D. C.

Von Damm, K. L., S. E. Oosting, R. Kozlowski, L. G.Buttermore, D. C. Colodner, H. N. Edmonds, J. M. Edmond,and J. M. Grebmeier (1995), Evolution of East Pacific Risehydrothermal vent fluids following a volcanic eruption,Nature, 375, 47–50, doi:10.1038/375047a0.

Walker, S. L., E. T. Baker, J. A. Resing, K. Nakamura, and P. D.McLain (2007), A new tool for detecting hydrothermalplumes: An ORP Sensor for the PMEL MAPR, Eos Trans.AGU, 88(52), Fall Meet. Suppl., Abstract V21D‐0753.

Welhan, J. A., and H. Craig (1983), Methane, hydrogen andhelium in hydrothermal fluids at 21°N on the East PacificRise, inHydrothermal Processes at Seafloor SpreadingCenters,edited by P. A. Rona et al., pp. 391–410, Plenum, New York.

Wilcock,W. S. D. (1997), Amodel for the formation of transientevent plumes above mid‐ocean ridge hydrothermal systems,J. Geophys. Res., 102(B6), 12,109–12,121, doi:10.1029/97JB00512.

Young, C., and J. E. Lupton (1983), An ultra tight fluid sam-pling system using cold‐welded copper tubing, Eos Trans.AGU, 64, 735.

Zellmer, K. E., and B. Taylor (2001), A three‐plate kinematicmodel for Lau Basin opening, Geochem. Geophys. Geosyst.,2(5), 1020, doi:10.1029/2000GC000106.


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Unique event plumes from a 2008 eruption on the Northeast Lau Spreading Center

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