Article - MyGeologyPage - University of California, Davis

Article 

Volume 12, Number 12 

28 December 2011 

Q12017, doi:10.1029/2011GC003639 

ISSN: 1525‐2027 

Environmental and chemical controls on palagonitization 

Bruce D. Pauly, Peter Schiffman, and Robert A. Zierenberg 

Department of Geology, University of California, 1 Shields Avenue, Davis, 

California 95616, USA (bdpauly@ucdavis.edu) 

David A. Clague 

Monterey Bay Aquarium Research Institute, 7700 Sandholdt Road, Moss Landing, 

California 95039, USA 

[1] Palagonitized sideromelane from submarine volcaniclastic, seafloor volcanic, marine phreatomagmatic, 

lacustine phreatomagmatic, and subglacial volcanic settings was investigated using in situ microanalysis 

to test if palagonite composition and texture are related to depositional environment. Palagonitization extent 

varies linearly and inversely with original sample porosity, suggesting that porosity is a controlling factor of 

palagonitization. Water absorbance of reflected infrared light varies linearly with water content derived from 

electron microprobe totals. Palagonite water content has a linear, inverse relationship to palagonitization 

extent. REEs are immobile during palagonitization, so they can be used to construct isocon diagrams 

for estimating major‐element concentration changes. Major‐element and overall mass changes during 

palagonitization vary widely (particularly for FeO and TiO 2 ) and indicate that palagonitization cannot be 

an isovolumetric process. These parameters depend strongly on original sideromelane composition, thus 

requiring composition to be taken into account when performing global oceanic cation flux calculations. 

Subalkaline sideromelane dissolves much more rapidly than alkaline sideromelane during palagonitization. 

Two styles of palagonitization, burial‐diagenesis (relatively long‐duration, low water/rock; passive fluid 

circulation) and hydrothermal (relatively short‐duration, high water/rock; hydrothermal fluid circulation), 

are recognized. Observed palagonite REE concentration gradients indicate that sideromelane dissolution 

must continue in the zone behind the advancing palagonitization front. MgO was found to be highly mobile 

during palagonitization. Observed palagonite MgO gradients are not developed during sideromelane dissolution, 

but instead record initiation of syn‐ and/or post‐palagonitization conversion of the gel‐palagonite 

layer to a phyllosillicate layer, consistent with evolution of sideromelane alteration layers toward equilibrium 

with the solution. 

Components: 13,000 words, 14 figures, 7 tables. 

Keywords: alteration layers; equilibrium; glass dissolution; mass change; palagonitization; water content. 

Index Terms: 1039 Geochemistry: Alteration and weathering processes (3617); 8424 Volcanology: Hydrothermal systems 

(0450, 1034, 3017, 3616, 4832, 8135); 8427 Volcanology: Subaqueous volcanism. 

Received 28 June 2011; Revised 27 October 2011; Accepted 29 October 2011; Published 28 December 2011. 

Pauly, B. D., P. Schiffman, R. A. Zierenberg, and D. A. Clague (2011), Environmental and chemical controls on 

palagonitization, Geochem. Geophys. Geosyst., 12, Q12017, doi:10.1029/2011GC003639. 

Copyright 2011 by the American Geophysical Union 1 of 26

Geochemistry 

Geophysics 

Geosystems G 3 PAULY ET AL.: GEOCHEMICAL CONTROLS ON PALAGONITIZATION 10.1029/2011GC003639 

1. Introduction 

[2] Palagonitization entails the hydrolytic alteration 

of sideromelane whereby basaltic glass is dissolved 

and hydrated, producing palagonite and various 

authigenic minerals, notably zeolites and smectites 

[Fisher and Schmincke, 1984; Stroncik and 

Schmincke, 2001; Walton and Schiffman, 2003]. 

Palagonitization is a globally significant geochemical 

process. In spite of important implications to our 

understanding of nuclear waste storage [Crovisier 

et al., 2003], volcanic edifice stability [Schiffman 

et al., 2006], and Martian crustal evolution 

[Bishop et al., 2002], the controls and mechanism of 

the palagonitization reaction remain poorly understood. 

Palagonite generally is considered to be the 

initial, possibly metastable, replacement product of 

sideromelane during palagonitization, with the final 

replacement product being smectite [Staudigel and 

Hart, 1983; Thorseth et al., 1991; Stroncik and 

Schmincke, 2001]. In this study, we follow Walton 

and Schiffman [2003] and consider palagonitization 

to be the process of forming palagonitized 

sideromelane, which herein will be referred to simply 

as palagonite. We use the term palagonite to 

describe the gel‐like, yellow to orange (as viewed in 

plane‐polarized light), isotropic replacement product 

of sideromelane; this is the practical, comprehensive 

sense as defined by Stroncik and Schmincke 

[2001]. 

[3] Published major‐element compositional data on 

palagonites from different localities, examples of 

which can be found in the works by Honnorez [1967, 

1972], Hay and Iijima [1968a], Staudigel and Hart 

[1983], Furnes [1984], Bednarz and Schmincke 

[1989], Jercinovic et al. [1990], Crovisier et al. 

[1992], Daux et al. [1994], Stroncik and Schmincke 

[2001], and Walton and Schiffman [2003], show 

substantial variation. The cause(s) of these compositional 

variations remain poorly understood. Although 

there is general agreement among previous workers 

that dissolution (congruent or incongruent) and 

hydration are responsible for the creation of palagonite 

[Crovisier et al.,1987;Daux et al., 1994; Le Gal 

et al., 1999], the mechanism(s) by which the transformation 

occurs is not completely understood. 

Stroncik and Schmincke [2001] proposed that palagonitization 

consists of two different reaction stages, 

the first being congruent dissolution of glass and 

simultaneous precipitation of palagonite, the second 

being an aging process during which palagonite 

reacts with fluid and crystallizes to smectite. 

Crovisier et al. [2003] summarizes results also 

suggesting that palagonitization is not a steady state 

process because glass dissolution rates decrease by 

up to 5 orders of magnitude as the palagonite layer 

thickens. 

[4] The purpose of this study was to investigate 

whether palagonitization and the resulting textural 

and compositional properties of palagonites are 

controlled by characteristics of the starting materials 

and of the palagonitization environment. Selected 

palagonitized samples from various geologic settings 

and localities were systematically analyzed by 

optical petrography, electron microprobe analysis 

(EPMA), and laser ablation inductively coupled 

mass spectroscopy (LA‐ICPMS). An advantage of 

these techniques is the ability to perform in situ 

analyses of thin section samples of palagonite and 

adjacent sideromelane in their petrogenetic context. 

The widely used assumption that palagonite water 

content equals the difference between 100% and the 

total of the analyzed oxides in EPMA compositional 

analyses was tested using in situ reflected‐light 

Fourier transform infrared spectroscopy (RL‐FTIR). 

[5] Basaltic glass alteration textures currently are 

categorized as either biotic or abiotic; Staudigel et al. 

[2008] provided a recent review. Thorseth et al. 

[1992] suggested a mechanism for the formation of 

bioalteration features in volcanic glass whereby 

microbes dissolve glass by changing the local contact 

area pH. Biotic versus abiotic glass alteration can be 

differentiated based on petrographically observable 

microtextures [Furnes et al., 2002, 2007]. Bioalteration 

textures generally fall into two categories: granular 

and tubular (tunneling) [Furnes et al., 2001; 

Cockell and Herrera, 2008]. Staudigel et al. [2008] 

provide detailed documentation of these textures. 

Many recent studies [Furnes et al.,2001;Walton and 

Schiffman, 2003; Staudigel and Furnes, 2004; 

Staudigel et al., 2004, 2006; Walton, 2008] document 

evidence that microorganisms play a role in palagonitization, 

particularly in the enhancement of sideromelane 

dissolution and involvement in precipitation 

of authigenic phases. Because this study focuses on 

abiotic palagonitization, samples with visual biotically 

mediated textures were excluded. 

2. Sample Selection 

[6] The sampling strategy for this study was designed 

to test if observed variations in the petrographic and 

chemical data are related to varying protolith and/ 

or palagonitization environment. Localities were 

chosen based on prior knowledge of the geologic 

history and therefore likely conditions attending 

sample palagonitization. All of the samples have 

2of26

Geochemistry 

Geophysics 


undergone palgonitization, and were classified by 

similarity of geologic setting into 5 environmental 

groups. Samples in the Submarine Volcaniclastic 

group consist of hyaloclastite that was deposited and 

subsequently buried in marine environments. The 

Seafloor Volcanic samples are from surficial hyaloclastite 

deposits or surficial lava flows located near 

the submarine edifice where they formed. Samples in 

the Marine Phreatomagmatic and Lacustrine Phreatomagmatic 

groups consist of hyaloclastite formed at 

tuff rings and tuff cones. The Subglacial Volcanic 

samples are from hyaloclastite deposits formed during 

eruptions of volcanoes beneath Icelandic glaciers. 

Table 1 lists the location, nomenclature, geologic 

setting, type, age, and reference(s) for each of the 

studied samples by environment group. The samples 

in this study are of Neogene and Quaternary age, the 

youngest being the two Surtsey samples, which were 

deposited between 1963 and 1968. 

3. Analytical Methods 

3.1. Petrography 

[7] One thin section of each sample was pointcounted 

using a mechanical stage with 1 mm gridspacing. 

The number of points counted per sample 

varied as a function of the available thin‐section 

area and ranged from 240 to 464. Point‐counting 

categories were sideromelane, lithic and crystal 

clasts, palagonite and smectite, undifferentiated 

zeolite, pore space, and “other” as necessary (clastic 

matrix, non‐replacive smectite, etc.). Minus‐cement 

porosity [Walton and Schiffman, 2003], an estimate 

of initial (depositional) porosity, was determined for 

each sample. For this study, minus‐cement porosity 

was defined as % pore space + % zeolite − % lithic 

and crystal clasts. This parameter can range from 

0 to 100 percent; lower values indicate lower initial 

porosity and vice versa. Palagonitization extent, 

defined here as % zeolite + % palagonite + % 

smectite (calculated as lithic and crystal clast and 

pore space free) was also calculated for each sample. 

This parameter can range from 0 to 100 percent; 

higher values indicate a greater extent of palagonitization 

and vice versa. 

3.2. Reflected Light Infrared Spectroscopy 

(RL‐FTIR) 

[8] Four thin‐section samples, with widely differing 

inferred palagonite water contents (based 

on electron microprobe compositional data) were 

selected for analysis by FTIR‐reflectance methods 

[Grzechnik et al., 1996; Moore et al., 2000] using a 

Nicolet Magna 750 FTIR spectrometer, fitted with 

a SpectraTech Analytical‐IR microscope, at the 

U.S. Geological Survey in Menlo Park, CA. The 

reflectance spectrum for each sample consisted of 

250 scans performed over the 6000–650 cm −1 

range using a 40 mm aperture. Background reflectance 

was measured using these same parameters 

and a polished gold target. The spectra have a 

resolution of 32 cm −1 and were corrected to provide 

a flat baseline, followed by a Kramers‐Kronig 

transformation between 4000 and 2000 wave numbers. 

All data collection and transformations were 

undertaken with Omnic® software provided with 

the Nicolet FTIR spectrometer. The Kramers‐ 

Kronig absorbance due to total water was determined 

by measuring the peak height at 3500 cm −1 , 

but this indicates relative sample water content only, 

since RL‐FTIR calibration standards for palagonite 

water content are not available. 

3.3. Electron Probe Microanalysis (EPMA) 

[9] Palagonite and sideromelane major‐element 

composition was analyzed using a Cameca SX‐100 

electron microprobe at the University of California, 

Davis. Prior to analysis, a ∼250 Å‐thick conductive 

carbon coating was applied to the sample thin sections. 

All analyses were performed using 15 KeV 

accelerating voltage. Sideromelane was analyzed 

using a 10 mm rastered beam and 10 nA current. 

Elemental thermal stability (particularly of Na) in 

palagonites as a function of current and raster size 

was investigated; optimal beam settings for palagonite 

analyses were 5 mm raster size and 2 nA current. 

Peaks and backgrounds for all analyses were typically 

counted for 10 s, and standard ZAF correction 

techniques [Schiffman and Roeske, 2002] were used 

to convert net intensities relative to oxide and silicate 

calibration standards to concentrations. Two types 

of analyses were performed. Palagonite of all 

samples was initially analyzed at several spots 

roughly in the centers of the rinds; adjacent sideromelane 

spot analyses were also performed. Also, 

for several of the samples with rinds of sufficient 

thickness, spot analyses were performed at a uniform 

spacing along a traverse starting at the outer 

edge of the palagonitized rind, moving inward toward 

and perpendicular to the sideromelane contact, continuing 

across the contact, and ending with at least 

one analysis in sideromelane. The number of palagonite 

spot analyses possible for each traverse was a 

function of the rastered beam size and the particular 

sample’s palagonitized rind thickness. 


Geochemistry 

Geophysics 


Table 1. Sample Geologic Information and Reference(s) 

ENVIRONMENT GROUP 

Sample (Nomenclature for This Sudy) [Location] Geologic Setting Sample Setting (Rock Type) Age Reference(s) 

SUBMARINE VOLCANICLASTIC 

Hawaii Scientific Drilling Project 

(HSDP) 

[Big Island of Hawaii] (hyaloclastite) 

submerged volcano flank drillcore Pleistocene Walton and Schiffman [2003]; 

Garcia et al. [2007] 

St. of Hawaii Geothermal Hole #1 

(SOH‐1) 

[Big Island of Hawaii] (hyaloclastite) 

submerged volcano flank drillcore Pleistocene Novak and Evans [1991]; 

Trusdell et al. [1999] 

Palagonia submerged lava flows surface (uplifted) Miocene–Early Pliocene Honnorez [1967, 1972] 

[Acqua Amara, Palagonia, Sicily] (hyaloclastite) 

Hilina Bench volcaniclastic basin submarine surface Pleistocene Lipman et al. [2002, 2006]; 

Coombs et al. [2006]; 

Kimura et al. [2006] 

[Offshore, Big Island of Hawaii] (volcaniclastic hyaloclastite) 

SEAFLOOR VOLCANIC 

Davidson seamount submarine push-core 14.8–9.8 Ma Davis et al. [2002]; 

Davis and Clague [2003]; 

Clague et al. [2009] 

[Davidson Seamount, Pacific Ocean] (flow breccia) 

North Arch volcanic field submarine surface Pliocene–Pleistocene Clague et al. [1990, 2002]; 

Davis and Clague [2006] 

[North Arch volcanic field, 

offshore Hawaii] 

(flow margin; Mn oxide‐encrusted) 

O’ahu cone submarine surface 510 Ka Clague et al. [2006] 

[O’ahu, northeast offshore flank] (hyaloclastite) 

MARINE PHREATOMAGMATIC 

Diamond Head tuff ring subaerial surface 360 Ka Winchell [1947]; 

Clague et al. [2006] 

[O’ahu] (hyaloclastite) 

Surtsey (less palagonitized sample) tuff cone subaerial surface 1963 AD Thorarinsson et al. [1964]; 

Jakobsson [1972, 1978]; 

[offshore Iceland] (hyaloclastite) Jakobsson and Moore [1986]; 

Jakobsson et al. [2000] 


Geochemistry 

Geophysics 


Table 1. (continued) 

ENVIRONMENT GROUP 

Sample (Nomenclature for This Sudy) [Location] Geologic Setting Sample Setting (Rock Type) Age Reference(s) 

Surtsey (more palagonitized sample) tuff cone subaerial surface 1963 AD Thorarinsson et al. [1964]; 

Jakobsson [1972, 1978]; 

[offshore Iceland] (hyaloclastite) Jakobsson and Moore [1986]; 

Jakobsson et al. [2000] 

Heimaklettur tuff cone subaerial surface approx. Nine Ka Mattsson and Höskuldsson [2003] 

[Westman Islands, Iceland] (hyaloclastite) 

LACUSTRINE PHREATOMAGMATIC 

Black Point tuff cone subaerial surface 13.61 Ka Benson et al. [1998] 

[Mono Lake, eastern California] (hyaloclastite) 

Pahvant Butte tuff cone subaerial surface 15.3 Ka Oviatt and Nash [1989]; 

White [1996, 2001] 

[South‐central Utah] (hyaloclastite) 

Fort Rock tuff ring subaerial surface 100–50 Ka Heiken et al. [1981]; 

Jensen [1995] 

[Southeastern Oregon] (hyaloclastite) 

SUBGLACIAL VOLCANIC 

Mosfell (less palagonitized sample) tuya subaerial surface 400–500 Ka H. Franzson (personal 

communication, 2010) 

[Iceland] (Grimsnes district, Iceland) (hyaloclastite) 

Mosfell (more palagonitized sample) tuya subaerial surface 400–500 Ka H. Franzson (personal 

communication, 2010) 

[Iceland] (Grimsnes district, Iceland) (hyaloclaastite) 

Ingólfsfjall tuya subaerial surface approx. One Ma Saemundsson et al. [2010] 

[Iceland] (hyaloclastite) 


Geochemistry 

Geophysics 


Figure 1. Thin section photomicrograph (PPL) of HSDP sample R0777–17.4 (hyaloclastite; submarine volcaniclastic 

group). Note relatively thick, smooth, gel‐like palagonite rinds developed on sideromelane grains, abundant zeolites 

(phillipsite and chabazite), and evidence of original porosity reduction due to burial such as abundant fractured 

grains and intergranular contacts. 

3.4. Laser Ablation Mass Spectroscopy 

(LA‐ICPMS) 

[10] Trace element analyses were performed using 

Laser Ablation‐Inductively Coupled Mass Spectrometry 

(LA‐ICPMS) at the University of 

California‐Davis Interdisciplinary Center for Plasma 

Mass Spectrometry. A New Wave Research Nd: 

YAG 213 nm laser (model UP‐213), equipped with 

an optical microscope, was used for laser ablation. 

Helium gas was used to carry ablated material to an 

Agilent Technologies inductively coupled mass 

spectrometer (model 7500a) equipped with a quadrupole 

mass analyzer and an electron multiplier 

detector. Ar gas was mixed with the He carrier gas at 

the entrance to the plasma chamber. Palagonitized 

sideromelane can disintegrate easily with application 

of the laser. In order to prevent uncontrolled ablation, 

optimal laser parameters were determined by trial and 

error, attempting to maximize count rates and prevent 

sample destruction during analysis. Sample thin 

sections were ablated using a 20‐micron beam size, 

70% laser energy (approximately 3.75 J/cm 2 ), and a 

10 Hz laser pulse repetition rate. For each analysis, 

the laser dwell time was 35 s. Mass spectrometer 

integration times were 10 ms for V, Sr, Zr, and Rb, 

30 ms for Sc, Cr, Co, and Cu, and 50 ms for Zn, Nb, 

Ba, REE, and Pb. U.S. Geological Survey Geochemical 

Reference Materials BCR‐2 (absolute external 

calibration) and BHVO‐2 (internal reference) fused to 

glasses were used as calibration standards. Intensities 

(counts) were converted to concentrations by the 

GLITTER software [van Achterbergh et al., 2001]. 

Palagonite and adjacent sideromelane spot and traverse 

(when possible) analyses locations were as 

described above for EPMA. 

4. Results 

4.1. Petrography 

[11] Significant palagonite microtextural differences 

across the sample set were observed with respect to 

the abundance of pore space (initial and remaining), 

abundance of zeolitic cements, and thickness of 

palagonitized rinds. Submarine volcaniclastic samples 

such as those from HSDP have relatively abundant 

zeolites and evidence of significant porosity 

reduction (likely due to their deep burial) such as 

significant fractures and grain to grain contacts 

(Figure 1). These samples also have relatively thick 

and smooth palagonitized rinds on sideromelane 

grains. In contrast, samples from the seafloor volcanic, 

marine phreatomagmatic (example Surtsey), and 

lacustrine phreatomagmatic groups typically have 

relatively rare zeolites, abundant pore space, and thin 

palagonitized rinds (Figure 2). The subglacial volcanic 

samples have more variability in terms of zeolite 

abundance, porosity, and palagonitized rind 

thickness. Measured palagonitized rind thickness 

varies widely across the sample set from 0.020 to 

0.600 mm (Table 2). 


Geochemistry 

Geophysics 


Figure 2. Thin section photomicrograph (PPL) of Surtsey sample 07010405 (the more palagonitized of the 2 Surtsey 

samples; hyaloclastite; marine phreatomagmatic group). Note relatively thin palagonitized rinds, rare zeolites, and 

abundant pore space. 

Table 2. 

Rind Thickness, Minus‐Cement Porosity, and Palagonitization Extent Results 

Environment Group 

Sample Name 

Sample 

Identification 

Maximum Rind 

Thickness (mm) 

Percent Palagonite 

and Assoc. 

Smectite 

Percent 

Zeolite 

Minus‐Cement 

Porosity a (%) 

Palagonitization 

Extent b (%) 

Submarine Volcaniclastic 

HSDP R0777–17.4 0.350 28 21 37 85 

SOH‐1 4565 0.200 49 17 19 71 

Hilina Bench (all grains) S708–9 0.225 47 26 36 90 

Hilina Bench (subalkaline grains) S508–7 0.225 

Hilina Bench (alkaline grains) S508–7 0.150 

Palagonia PAL‐1 0.425 55 20 26 88 

Seafloor Volcanic 

Davidson T430‐R8 – c 31 16 43 70 

North Arch 26D‐6 0.200 37 3 51 57 

O’ahu T298‐R17 0.250 32 5 56 53 

Marine Phreatomagmatic 

Diamond Head 02010405 0.125 32 19 59 66 

Surtsey (less palagonitized) 07010404 – d 9 9 71 29 

Surtsey (more palagonitized) 07010405 0.150 26 11 63 58 

Heimaklettur 07020402 – d 40 12 30 68 

Lacustrine Phreatomagmatic 

Black Point 08310202 0.020 12 12 48 40 

Pahvant Butte 10230304 0.300 35 28 51 84 

Fort Rock 08029904 0.600 42 19 50 76 

Subglacial Volcanic 

Mosfell (less palagonitized) MOS‐1 0.040 14 9 24 35 

Mosfell (more palagonitized) MOS‐2 0.060 36 8 16 50 

Ingólfsfjall ING‐1 0.325 52 26 34 82 

a Minus‐cement porosity (%) = % pore space + % zeolites − % lithic and crystal clasts. 

b Palagonitization extent (%) = % zeolites + % palagonite and smectite (calculated as lithic and crystal clast and pore space free). 

c Rind thickness not measurable (glassy flow margin; palagonitized in areas, not rinds). 

d Rind thickness not measurable (rinds are very thin and incomplete). 


Geochemistry 

Geophysics 


Figure 3. Plot of palagonitization extent versus minus‐cement porosity. Black solid symbols: submarine volcaniclastic 

group; violet solid symbols: seafloor volcanic group; blue solid symbols: marine phreatomagmatic group; 

red open symbols: lacustrine phreatomagmatic group; green symbols: subglacial volcanic group. For the submarine 

volcaniclastic, seafloor volcanic, and marine phreatomagmatic environment groups, samples with relatively greater 

palagonitization extent also have relatively low minus‐cement porosity and vice versa; samples from the other two 

groups are more variable. However, the more‐palagonitized Surtsey and Mosfell samples have higher palagonitization 

extent and lower minus‐cement porosity than the corresponding less‐palagonitized samples from the same edifices 

(colored arrows connecting symbols), implying that samples with lower original porosity (and expected lower permeability) 

palagonitized to a greater extent. 

[12] Minus‐cement porosity and palagonitization 

extent (Table 2) also show wide variations. Minuscement 

porosity ranges from 16 (relatively low) to 

71% (relatively high initial pore space). Minuscement 

porosity is fairly consistent within each 

environmental group (except for the Heimaklettur 

sample of the marine phreatomagmatic group) and is 

relatively low for the submarine volcaniclastic and 

subglacial volcanic samples. Palagonitization extent 

ranges from 29 (least palagonitized) to 90% (most 

thoroughly palagonitized). Palagonitization extent 

is consistently higher for the submarine volcaniclastic 

samples; it is generally lower and more variable 

for samples from the other sample groups. 

Palagonitization extent versus minus‐cement 

porosity is plotted in Figure 3. The submarine volcaniclastic 

samples have relatively high palagonitization 

extent and relatively low minus‐cement 

porosity. The other groups are more variable, but 

in general, there is a trend of increasing palagonitization 

extent with decreasing minus‐cement porosity. 

The less‐palagonitized samples from Surtsey and 

Mosfell have greater minus‐cement porosities than the 

more‐palagonitized samples from these same edifices. 

4.2. Water Content Measurement 

[13] The Kramers‐Kronig absorbances measured 

using RL‐FTIR at 3500 cm −1 (the wave number of 

the total H 2 O absorbance peak) correspond to the 

derived (EPMA) water contents (Table 3). Figure 4 

shows that this is a nearly linear relationship, with 

the line of best fit having the equation 

Kramers-Kronig absorbance 

¼ 0:0081ðwt:% EPMA-derived H 2 OÞ: 

[14] The extinction coefficient of 0.0081 is in close 

agreement with the calibrated extinction coefficient 

obtained for basalt glass (0.0084) by King et al. 

[2006] using the same method. These workers calibrated 

the extinction coefficient using transmission 

IR spectroscopy on samples of natural sideromelane 

from the Manus Basin and the Galapagos Islands. 


Geochemistry 

Geophysics 


Table 3. Measured Kramers‐Kronig Absorbance for the 3500 cm −1 Band (RL‐FTIR) and Inferred Water Content (EPMA, 

wt. % Oxide) 

Location Sample KK Absorbance (RL‐FTIR) Inferred H 2 O (EPMA) 

Surtsey (less palagonitized sample) 07010404 0.260 33.4 

Davidson 26D‐6 0.246 27.1 

HSDP 7231 0.154 20.4 

Palagonia PAL‐1 0.130 14.0 

The precise extinction coefficient for palagonite 

awaits analysis of a set of palagonite samples with 

known absolute water contents. The linear relationship 

observed here, however, supports our 

assertion that the inferred water content (i.e., 100% 

minus the EPMA analysis total) of a palagonite 

sample represents its water content, and that therefore 

palagonite water content can be derived with the 

other major‐elements using EPMA. Water content 

results are given and discussed below (majorelement 

composition). 

4.3. Major‐Element Composition 

[15] Sideromelane major‐element composition results 

(EPMA) are given in Table 4. Sampled sideromelane 

compositions are fairly evenly distributed between 

alkaline and subalkaline (Figure 5). Averaged palagonite 

compositions for each sampled locality (measured 

adjacent to the sideromelane analyses in the 

center of the palagonitized rind) are given in Table 5. 

[16] Sample S508–7 from Hilina Bench is unique in 

this study because it contains both alkaline and subalkaline 

palagonitized sideromelane grains (Figure 6). 

These grains are from the pre‐shield alkalic and shield 

tholeiitic stages [Sisson et al., 2002, and references 

therein], respectively, of Kilauea volcano development. 

Palagonitization of these samples must have 

occurred after final deposition within the (Hilina) 

volcaniclastic basin, because the hyaloclastite grains 

are for the most part sub‐rounded and have relatively 

thick and smooth palagonite rinds, which would not 

be expected to survive erosion. Palagonitized rinds 

for this and other samples were observed to be so 

friable that preparation of double‐sided thin sections 

was not possible without rind separation from the 

fresh glass. Therefore, all of the grains within this 

particular sample must have palagonitized with 

identical conditions and for identical duration. The 

major‐element compositions of sideromelane and 

corresponding palagonite for 2 adjacent pairs of 

Figure 4. Plot of measured Kramers‐Kronig absorbance (from RL‐FTIR measurements) versus inferred water content 

(from EPMA analyses) for samples from the 4 localities indicated (Table 3). A nearly linear, direct relationship 

between absorbance and inferred water content is indicated. 


Geochemistry 

Geophysics 


Table 4. 

Average Sideromelane Major‐Element Compositions (wt. % Oxide), Determined Using EPMA 

Sample Group 

Location 

(Sample ID) 

Number 

of 

Concentration (wt. % Oxide): 

Sideromelane 

Analyses SiO 2 TiO 2 Al 2 O 3 FeO MnO MgO CaO Na 2 O K 2 O 

P 2 O 5 SO 2 Total 


HSDP (R0777–17.4) 10 52.66 2.60 13.83 10.77 0.18 6.72 10.63 2.31 0.31 0.24 0.03 100.65 

SOH‐1 (3135) 9 50.76 2.86 13.63 11.36 0.16 6.09 10.28 2.54 0.48 0.38 0.07 98.60 

Palagonia (PAL‐1) 9 52.75 1.72 14.66 9.45 0.13 6.39 9.42 3.08 0.19 0.28 0.11 98.19 

Hilina Bench (S508‐7) (See Table 6) 

Submarine Volcanic 

Davidson (T430‐R8) 9 44.18 3.40 17.39 9.22 0.14 4.59 8.61 4.37 1.79 0.74 0.11 94.53 

North Arch (26D‐6) 3 44.64 2.21 13.79 11.44 0.24 6.73 13.77 3.82 0.91 0.35 0.35 98.25 

O’ahu (T298‐R17) 9 40.74 3.39 12.99 11.86 0.20 5.24 14.79 5.25 1.82 1.19 0.30 97.77 


Diamond Head (02010405) 9 42.65 2.70 13.84 11.87 0.18 6.38 13.42 3.94 0.98 0.66 0.12 96.74 

Surtsey–less palagonitized 9 46.33 2.37 15.70 12.01 0.22 5.82 10.25 3.84 0.55 0.53 0.12 97.74 

sample (07010404) 

Surtsey–more palagonitized 9 48.53 2.48 14.56 12.24 0.23 5.70 10.74 3.75 0.54 0.50 0.12 99.39 

sample (07010405) 

Heimaklettur (07020402) 9 46.32 2.66 15.78 12.18 0.21 5.87 10.02 3.71 0.70 0.39 0.10 97.93 


Black Point (09120301) 6 49.90 1.51 18.78 8.32 0.13 5.63 8.59 3.94 1.08 0.36 0.03 98.27 

Pahvant Butte (10230304) 5 51.76 1.75 15.17 12.01 0.18 4.87 8.36 3.40 1.13 0.39 0.09 99.12 

Fort Rock (08029904) 4 50.24 1.65 15.60 11.12 0.22 5.90 8.78 3.55 0.45 0.34 0.09 97.95 


Mosfell–less palagonitized 9 47.05 1.50 15.60 10.82 0.21 8.65 12.68 2.17 0.05 0.22 0.02 98.96 

sample (MOS‐1) 

Mosfell–more palagonitized 11 47.78 1.67 16.01 11.11 0.18 8.70 12.66 2.21 0.05 0.14 0.02 100.53 


Ingolfsfjall (ING‐1) 9 48.96 1.72 15.04 11.92 0.20 7.90 11.72 2.45 0.14 0.25 0.13 100.44 

Figure 5. Total alkalies versus silica plot of sideromelane compositions, averaged at each locality sampled. Black 

solid symbols: submarine volcaniclastic group; violet solid symbols: seafloor volcanic group; blue solid symbols: 

marine phreatomagmatic group; red open symbols: lacustrine phreatomagmatic group; green symbols: subglacial volcanic 

group. Sideromelane composition sampled in this study covers a broad range and is fairly evenly distributed 

across alkaline and subalkaline compositions. 

10 of 26

Geochemistry 

Geophysics 


Table 5. 

Average Palagonite Major‐Element Compositions (wt. % Oxide), Determined Using EPMA 

Sample Group 

Location 

(Sample ID) 

Number 

of 

Concentration (wt. % Oxide): 

Palagonite 

Analyses SiO 2 TiO 2 Al 2 O 3 FeO MnO MgO CaO Na 2 O K 2 O 

P 2 O 5 SO 2 Total H 2 O a 


HSDP (R0777‐17.4) 9 40.80 5.15 9.88 11.50 0.00 0.97 10.07 0.91 0.16 0.10 0.02 79.62 20.38 

SOH‐1 (3135) 9 41.87 4.70 9.31 15.45 0.16 2.73 9.23 0.91 0.35 0.37 0.11 85.19 14.81 

Palagonia (PAL‐1) 9 44.54 3.27 10.65 16.16 0.11 1.86 7.13 1.83 0.32 0.04 0.03 85.99 14.01 

Hilina Bench (S508‐7) (See Table 6) 

Seafloor Volcanic 

Davidson (T430‐R8) 8 38.25 3.12 10.47 12.32 0.00 1.09 0.79 4.57 1.94 0.21 0.15 72.91 27.09 

North Arch (26D‐6) 4 26.40 4.47 4.93 20.96 0.03 0.83 0.71 2.24 0.83 0.00 0.11 61.57 38.43 

O’ahu (T298‐R17) 9 31.41 4.87 16.08 16.84 0.00 2.06 2.40 3.07 1.13 1.70 0.14 79.70 20.30 


Diamond Head (02010405) 8 43.06 4.25 7.83 19.07 0.02 0.78 4.30 0.34 0.43 0.08 0.01 80.21 19.79 

Surtsey–less palagonitized 9 27.80 3.99 7.55 16.41 0.19 2.22 6.87 0.72 0.16 0.70 0.00 66.64 33.36 

sample (07010404) 

Surtsey–more palagonitized 9 30.39 3.81 12.84 18.46 0.16 3.31 6.89 0.04 0.04 0.30 0.00 76.39 23.61 

sample (07010405) 

Heimaklettur (07020402) 6 37.01 4.28 10.40 17.23 0.07 2.31 4.52 0.45 0.66 0.00 0.01 77.73 22.27 


Black Point (09120301) 9 36.99 2.92 9.49 12.98 0.06 1.12 6.76 1.85 0.78 0.20 0.01 73.20 26.80 

Pahvant Butte (10230304) 3 36.81 2.91 8.25 16.35 0.05 0.55 8.06 0.68 0.14 0.06 0.01 73.88 26.12 

Fort Rock (08029904) 5 34.07 2.27 10.08 15.46 0.01 1.05 4.66 1.27 0.25 0.11 0.01 69.28 30.72 


Mosfell–less palagonitized 9 42.58 2.58 11.15 17.35 0.07 0.75 9.83 0.17 0.05 0.08 0.03 84.66 15.34 


Mosfell–more palagonitized 3 46.81 2.79 13.54 13.12 0.05 2.92 6.59 0.19 0.05 0.05 0.03 86.34 13.66 


Ingólfsfjall (ING‐1) 9 42.33 3.33 11.00 17.23 0.14 3.08 8.34 0.14 0.16 0.11 0.09 85.99 14.01 

a Calculated as 100‐Total. 

alkaline and subalkaline palagonitized hyaloclastite 

grains (Figure 6) in this sample are given in Table 6. 

The alkaline versus subaklaline compositions of the 

sideromelane grains are clearly distinguishable 

(Table 6 and Figure 6 (insets)), as are those of the 

corresponding palagonites (Table 6). Also, it can be 

seen in Figure 6 that the palagonitized margins of the 

subalkaline grains are relatively thick (Table 2) and 

deep‐orange colored, whereas palagonitized margins 

of the alkaline grains are relatively thin and paleorange 

colored. 

[17] The derived palagonite water contents of all 

samples are given in Tables 5 and 6, and are plotted 

(by sample location and environmental group) in 

Figure 7. The averaged group values for water 

content range from approximately 14 (subglacial 

volcanic group) to 29 (seafloor volcanic group) 

wt. % oxide. The values are generally consistent 

within the seafloor volcanic, lacustrine phreatomagmatic, 

and subglacial volcanic groups, although 

the less‐palagonitized Surtsey sample has significantly 

higher water content than the other marine 

phreatomagmatic samples. The palagonitized rinds 

developed on the 2 subalkaline grains analyzed in 

Hilina Bench sample S508–7 (Figure 6 and Table 6) 

have much higher water content (25% average) than 

for the 2 alkaline grains (17% average). Water 

content (Tables 6 and 7) versus palagonitization 

extent (Table 2) is plotted in Figure 8. Water content 

generally decreases with increasing palagonitization 

extent. This trend also holds for the less‐ and morepalagonitized 

Surtsey and Mosfell samples, with 

both sample sets showing lower water content with 

higher palagonitization extent. No correlation 

between sample age and sample water content was 

observed. 

4.4. Trace Element Composition 

[18] The raw trace element data are presented in 

Table S1 in the auxiliary material. 1 Multielement 

spider diagrams were prepared for each sample 

1 Auxiliary materials are available in the HTML. doi:10.1029/ 

2011GC003639. 


Geochemistry 

Geophysics 


Figure 6. Thin section photomicrographs (PPL) and TAS plot (inset) for Hilina Bench sample S508–7. Sideromelane 

of hyaloclastite grains (a) 12 and 13 and (b) 18 and 19 are alkaline and subalkaline, respectively (Table 6). Note 

that the palagonitized rinds on the subalkaline grains are colored a deeper shade of orange than those of the alkaline 

grains (Table 2), which are relatively pale (the sideromelanes are also distinguishably colored). The palagonitized 

rinds developed on the subalkaline grains are also significantly thicker than those on the alkaline grains. 

Table 6. 

Compositional Analyses (EPMA) of Sideromelane and Palagonite Grains in Hilina Bench Sample S508–7 a 

SiO 2 TiO 2 Al 2 O3 FeO MnO 

Concentration 

(wt. % Oxide) 

MgO CaO Na 2 O 

K 2 O P 2 O 5 SO 2 Total H 2 O b 

Sideromelane 

S508‐7 G12‐1 (alkaline) 47.81 4.17 12.96 13.02 0.18 6.08 11.13 2.93 0.64 0.46 0.08 99.45 

S508‐7 G12‐2 (alkaline) 47.98 4.21 13.50 13.20 0.18 6.08 11.14 2.92 0.75 0.55 0.01 100.52 

S508‐7 G18‐1 (alkaline) 46.22 4.68 13.79 13.44 0.16 5.96 11.13 3.25 0.84 0.70 0.07 100.24 

S508‐7 G18‐2 (alkaline) 46.32 4.64 13.98 13.53 0.25 5.92 11.15 3.25 0.88 0.66 0.04 100.61 

S508‐7 G18‐3 (alkaline) 46.01 4.56 13.63 13.24 0.20 6.02 11.35 3.31 0.89 0.47 0.02 99.70 

average alkaline 46.87 4.45 13.57 13.28 0.19 6.01 11.18 3.13 0.80 0.57 0.04 100.10 

S508‐7 G13‐1 (subalk.) 52.29 2.30 13.21 11.44 0.18 6.84 10.96 2.27 0.32 0.41 0.02 100.25 

S508‐7 G13‐2 (subalk.) 52.24 2.33 13.51 11.46 0.20 6.72 10.93 2.38 0.31 0.49 0.00 100.55 

S508‐7 G13‐3 (subalk.) 52.57 2.29 13.40 11.20 0.20 6.74 10.98 2.26 0.33 0.37 0.00 100.34 

S508‐7 G19‐1 (subalk.) 53.70 2.20 13.47 10.98 0.20 7.00 10.82 2.36 0.31 0.36 0.00 101.40 

S508‐7 G19‐2 (subalk.) 52.95 2.14 13.73 10.82 0.18 7.13 10.91 2.29 0.29 0.32 0.03 100.78 

average subalkaline 52.75 2.25 13.46 11.18 0.19 6.89 10.92 2.31 0.31 0.39 0.01 100.66 

Palagonite 

S508‐7 G12‐1 (alkaline) 36.75 6.97 6.53 12.27 0.10 3.45 3.59 1.95 1.31 0.15 0.09 73.14 26.86 

S508‐7 G12‐2 (alkaline) 38.76 7.60 7.24 12.17 0.08 2.86 3.90 1.63 1.05 0.17 0.08 75.53 24.47 

S508‐7 G18‐1 (alkaline) 39.08 8.32 8.10 10.51 0.16 2.78 4.73 1.19 1.05 0.14 0.13 76.19 23.81 

S508‐7 G18‐2 (alkaline) 39.56 8.33 7.65 11.30 0.04 2.58 4.88 1.98 0.98 0.00 0.14 77.45 22.55 

S508‐7 G18‐3 (alkaline) 36.45 7.48 6.42 12.65 0.18 3.10 3.59 1.57 1.15 0.08 0.00 72.67 27.33 

average alkaline 38.12 7.74 7.19 11.78 0.11 2.95 4.14 1.66 1.11 0.11 0.09 74.99 25.0 

S508‐7 G13‐1 (subalk.) 39.72 4.57 8.03 18.79 0.11 3.14 2.62 1.96 1.46 0.00 0.32 80.71 19.29 

S508‐7 G13‐2 (subalk.) 42.01 4.74 9.60 17.23 0.18 2.67 3.39 1.93 1.45 0.15 0.19 83.53 16.47 

S508‐7 G13‐3 (subalk.) 42.09 4.79 8.40 17.66 0.07 3.79 2.76 2.20 1.37 0.11 0.07 83.29 16.71 

S508‐7 G19‐1 (subalk.) 43.29 4.51 10.22 15.50 0.14 2.87 3.66 2.27 1.45 0.00 0.02 83.91 16.09 

S508‐7 G19‐2 (subalk.) 42.46 4.35 9.57 15.96 0.07 2.62 3.33 2.44 1.56 0.01 0.08 82.45 17.55 

average subalkaline 41.91 4.59 9.16 17.03 0.11 3.02 3.15 2.16 1.46 0.05 0.14 82.78 17.22 

a Grains 12 and 18 consist of alkaline sideromelane (dark blue); grains 13 and 19 consist of subalkaline sideromelane (red). 

b Calculated as 100‐Total. 


Geochemistry 

Geophysics 


Figure 7. Plot of derived water content (wt. % oxide; EPMA) versus sample location. Black: submarine volcaniclastic 

group; violet: seafloor volcanic group; blue: marine phreatomagmatic group; red: lacustrine phreatomagmatic 

group; green: subglacial volcanic group. Values within the seafloor volcanic and subglacial volcanic groups are fairly 

consistent; each of the other groups contains at least one sample with an inconsistent value. Average group values 

(indicated by the black lines) range from approximately 14 (subglacial volcanic group) to 29 wt. % oxide (seafloor 

volcanic group). 

Table 7. Major‐Element and Overall Mass Change (%) for Subalkaline and Alkaline Samples, Determined Graphically Using 

Isocon Diagrams (Figure 10 and Figure S2) a 

Location 

Mass Change (%) Element (Oxide) 

SiO 2 TiO 2 Al 2 O 3 FeO MgO CaO Na 2 O K 2 O 

Overall Mass 

Change (%) 

(Subalkaline Sideromelane) 

HSDP* −71.7 −31.1 −70.9 −68.4 −96.1 −63.9 −87.4 −77.2 −63.7 

SOH‐1* −54.3 6.3 −64.2 −21.8 −63.7 −41.6 −82.5 −49.5 −42.6 

Palagonia* −56.3 2.8 −66.5 2.2 −82.7 −59.1 −69.1 −41.4 −47.0 

Ingólfsfjall −58.9 −5.4 −62.8 −41.7 −84.3 −61.8 −97.9 −38.9 −53.3 

Hilina Bench (subalk. grain)* −48.7 34.0 −59.8 1.0 −64.0 −83.9 −37.6 166.0 −35.9 

Avg. Subalkaline Sideromelane −58.0 1.3 −64.8 −25.7 −78.1 −62.1 −74.9 −8.2 −48.5 

(Alkaline Sideromelane) 

Davidson* −53.0 −47.6 −70.3 −17.5 −88.2 −95.0 −33.6 8.1 −43.2 

Surtsey (more palagonitized sample) −56.3 18.5 −33.8 8.7 −64.8 −49.8 −91.8 −91.8 −27.4 

O’ahu* −50.8 −8.1 −1.5 3.7 −83.0 −90.6 −69.4 −65.2 −33.0 

Black Point −40.8 109.5 −64.6 44.2 −71.3 −45.6 −69.5 −22.5 −23.4 

Pahvant Butte −53.0 7.8 −63.2 −5.5 −91.9 −27.8 −89.2 −92.8 −36.3 

Diamond Head −17.9 11.1 −57.7 32.3 −88.6 −79.9 −91.9 −65.3 −19.6 

Fort Rock −53.8 2.4 −53.7 3.8 −88.6 −71.0 −44.0 −60.7 −29.0 

Heimaklettur −55.9 27.6 −66.9 19.7 −83.9 −76.9 −63.5 −69.1 −38.1 

Hilina Bench (alkaline grain)* −26.4 54.7 −48.9 −31.9 −59.4 −63.0 −68.1 8.8 −13.0 

Avg. Alkaline Sideromelane −45.3 19.5 −51.2 6.4 −80.0 −66.6 −69.0 −50.0 −29.2 

*Oceanic sample average (this study) −51.6 1.6 −54.6 −19.0 −76.7 −71.0 −64.0 −7.2 

Staudigel and Hart [1983, Table 2]; MORB −50.0 n/a −55.0 −12.0 −67.0 −88.0 −81.0 900.0 

a The averaged values for the oceanic samples are compared to MORB values calculated by Staudigel and Hart [1983] at the bottom of the table 

(see section 5). 


Geochemistry 

Geophysics 


Figure 8. Plot of palagonite water content versus palagonitization extent. Black solid symbols: submarine volcaniclastic 

group; violet solid symbols: seafloor volcanic group; blue solid symbols: marine phreatomagmatic group; red 

open symbols: lacustrine phreatomagmatic group; green symbols: subglacial volcanic group. Samples with lower 

water contents generally palagonitized to a greater extent, however this trend is not evident in the seafloor volcanic 

and lacustrine phreatomagmatic groups. The less‐ and more‐palagonitized sample pairs from Surtsey and Mosfell 

show this trend (indicated by arrows connecting symbols). 

(Figures 9 and S1). The sampled palagonite REE 

patterns among the different environment groups 

and for localities within each environment group 

are broadly similar. Sr, K, Rb, and Ba in palagonite 

relative to sideromelane are highly variable across 

the sample set. Sr and K in palagonite vary from 

depleted to enriched (relative to sideromelane), 

independent of environment sample group. Rb and 

Ba in palagonite, for the majority of the samples, 

are depleted (in some cases greatly) relative to 

sideromelane. For most of the samples, Nb through 

Lu (including REEs) palagonite concentration 

Figure 9. Chondrite‐normalized [McDonough and Sun, 1995] multielement spider diagrams for (a) a subalkaline 

grain and (b) an alkaline grain in Hilina Bench sample S508–7 (submarine volcaniclastic environment group). Colors 

and analyses numbers as in Figure S1. Note the difference in Nb through Lu (immobile elements) concentration 

relative to sideromelane for the palagonitized subalkaline versus alkaline grain. 


Geochemistry 

Geophysics 


Figure 10. Isocon diagrams [Grant, 1986] used for determination of elemental and overall mass change for palagonite 

developed on (a) a subalkaline grain (plotted trace element analyses Sid. One and average of Pal. One and Pal. 2) and 

(b) an alkaline grain (plotted trace element analyses Sid. One and average of Pal. One and Pal. 2) in Hilina Bench sample 

S508–7 (submarine volcaniclastic environment group). For both samples, plotted major‐element analyses and trace 

element analyses were carried out using EPMA and LA‐ICPMS, respectively, at identical spots in sideromelane and 

palagonite. Plotted palagonite major‐ and trace element analyses were made as close to the center of palagonitized rinds 

as possible. Plotted sideromelane major‐ and trace element analyses were made at identical spots in sideromelane, which 

in all samples is homogeneous with respect to major‐ and trace element composition. Plotted major‐element sideromelane 

and palagonite analyses (green symbols) are given in Table S2; plotted trace element analyses (red and blue 

symbols) are given in Table S1. C A : factored concentration in altered sideromelane (palagonite); C O : original sideromelane 

factored concentration (see text). Isocon line (red) was determined by linear regression using Zr, Nb, Sc, La, 

and Nd. The slope of the isocon line is significantly steeper for palagonitized subalkaline (35.9% overall mass loss) than 

for alkaline sideromelane (13.0% overall mass loss). 

patterns are parallel to and show higher concentrations 

than the corresponding sideromelane patterns. 

Walton et al. [2005], using LA‐ICPMS, found 

similar behavior in palagonite relative to sideromelane 

REE patterns for HSDP hyaloclastite 

samples, as did Daux et al. [1994], using inductively 

coupled plasma atomic emission spectrometry 

(ICP‐AES) to analyze chemically separated 

sideromelane and palagonite fragments from Icelandic 

subglacially erupted hyaloclastite samples. 

The immobile‐element sections of the spider patterns 

of the present study indicate that essentially 

no systematic chemical fractionation of these 

elements occurred during palagonitization (a few 

samples show the LREEs to be slightly less enriched 

than the HREEs). This finding is consistent 

with Daux et al. [1994], who found no evidence for 

REE fractionation among a suite of 11 Icelandic 

palagonitized hyaloclastites. For samples with 

sufficiently thick palagonitized rinds, so that palagonite 

analyses could be determined along a traverse 

(Pal. 1, Pal. 2, etc. in Figures 9 and S1), 

immobile‐element concentration decreases systematically 

from the outer palagonitized rim inward 

toward the fresh sideromelane. Immobile‐element 

(Nb through Lu) concentration relative to sideromelane 

is greater for subalkaline (Figures S1a, 

S1b, S1c, and S1m) as compared to alkaline 

(Figures S1d, S1f, S1g, S1h, S1i, S1j, S1k, and 

S1l) palagonite. This is also the case for subalkaline 

and alkaline grains analyzed in Hilina Bench 

sample S508–7 (Figure 9). Only the North Arch 

palagonite trace element pattern (Figure S1e) does 

not follow the trends described above, and shows 

a distinct Ce anomaly. 

4.5. Major‐Element Concentration 

Changes 

[19] Isocon diagrams [Grant, 1986] were prepared for 

each sample in order to investigate major‐element 

concentration changes during palagonitization. This 

method allowed a test of a constant volume or constant 

mass assumption based on immobile element 

concentrations. The trace‐ and major‐element concentration 

analyses (Tables S1 and S2, respectively) 

used for these diagrams were from LA‐ICPMS and 

EPMA analyses performed at approximately the 

center of the palagonitized outer margins and at as 

close to identical spots as possible for each sample. 


Geochemistry 

Geophysics 


Figure 11. Mass loss (Table 7) versus palagonitization extent (Table 2). Samples plot along a general trend of 

increasing mass loss with increasing extent (arrow). Subalkaline samples for the most part have greater mass loss 

values than alkaline samples. The alkaline and subalkaline fields overlap for extent values between 70 and 85%, 

however in this region, subalkaline mass loss values are generally greater. 

The palagonite compositions were plotted against the 

analyses of adjacent sideromelane. Major‐element 

(wt. % oxide) and trace element (ppm) concentrations 

were scaled, using appropriate factors, for 

plotting. The actual trace‐ and major‐element analyses 

used to construct each sample’s isocon diagram 

are given with Figures 10 and S2. Zr, Nb, Sc, La, and 

Nd concentrations consistently plotted along isocon 

lines for almost all of the samples (Figures 10 and 

S2). These elements therefore can be considered 

immobile. Major‐element oxide as well as Rb, V, 

and Cr concentrations did not consistently plot 

along the isocon lines, hence these elements are 

considered mobile. Concentration changes for mobile 

elements, as well as overall mass loss, were determined 

graphically using each sample’s isocon diagram 

and the lever rule [Grant, 1986]. 

[20] The mass loss results fall into two categories: 

palagonites in samples of subalkaline sideromelane 

(Figures 10a, S2a, S2b, S2c, and S2m) exhibit relatively 

higher (35.9 to 63.7%) mass loss; palagonites 

in samples of alkaline sideromelane (Figures 10b, 

S2d, S2f, S2g, S2h, S2i, S2j, S2k, and S2l) exhibit 

relatively lower (13.0 to 43.2%) mass loss (Table 7). 

In particular, the palagonites in samples of subalkaline 

sideromelane lost significant amounts Si, 

Ti, Al, and Fe, whereas in palagonites of samples of 

alkaline sideromelane, relatively less Si and Al were 

lost, and, on average, Ti and Fe were gained. These 

trends are also observed for the subalkaline and 

alkaline grains in Hilina Bench sample S508–7 

(Figure 10 and Table 7). Overall mass loss versus 

palagonitization extent is plotted in Figure 11. Mass 

loss generally increases linearly with increasing 

palagonitization extent. Samples of subalkaline 

sideromelane that palagonitized to a relatively high 

extent also had relatively high mass loss. For a given 

palagonitization extent, samples of subalkaline palagonitized 

sideromelane lost more mass than alkaline 

samples. The North Arch Zr, Nb, Sc, La, and Nd 

concentrations do not plot along a line (Figure S2e), 

so the isocon approach cannot be used to determine 

valid palagonitization mass changes for this sample. 

The North Arch sample was the only Mn oxideencrusted 

sample studied (Table 1). Previous studies 

indicate that Mn and Fe oxides comprising marine 

ferromanganese crusts scavenge trace metals 

[Koschinsky and Hein,2003;Hein et al., 1997]. The 

anomalous mobile behavior of Zr, Nb, Sc, La, and 

Nd in the North Arch sample we interpret to be due 

to variable sequestering of these elements within 

the Mn oxide crust. Furthermore, we interpret the 

observed negative Ce anomaly to be the result of 

significant oxidative accumulation of Ce within 

the Mn oxide crust, as observed in ferromanganese 

crusts from the Central Pacific by Bau and 

Koschinsky [2009]. 


Geochemistry 

Geophysics 


Figure 12. Major‐element concentration traverse, SOH‐1 (submarine volcaniclastic environment group; subalkaline 

sideromelane). (top) BSE photomicrograph showing the palagonite traverse in red and the sideromelane traverse in 

brown. Spot‐spacing is given on the lower image. (bottom) Palagonite major‐element concentration (Table S3), 

plotted as a percentage of the outermost (Spot 1) palagonite concentration, for each spot along the palagonite traverse. 

4.6. Major‐Element Concentration 

Gradients 

[21] Major‐element spot analyses along traverses 

through palagonite were performed on 5 of the samples 

with relatively thick palgonitized rinds. The raw 

data are given in Table S3, and the traverse details and 

plotted results are given in Figures 12 and S3a–S3d. 

In each sample, the sideromelane is homogeneous 

with respect to major‐element composition. For all 

5 of the palagonites, a distinct MgO concentration 

gradient, highest at the outer palagonitized margin 

and decreasing inward toward the sideromelane 

contact to only about 50% of the outer margin value, 

was observed. Over the same interval, SiO2, Al2O3, 

and CaO concentrations generally increased by 5 to 

20%, whereas TiO2 and FeO concentrations generally 

remained steady or decreased. 

5. Discussion 

5.1. Palagonitization Extent 

[22] Previous workers have devised various means 

for the quantification of palagonitization extent, 

based for example on palagonitized rind thickness 

and major‐element concentration changes during 

palagonitization. Interpretation of palagonitized rind 

thickness is not straightforward. Rind thickness has 

been shown to relate to both palagonitization duration 

[Moore, 1966] and temperature [Moore et al., 


Geochemistry 

Geophysics 


Figure 13. Palagonitization extents determined using 3 different methods (see legend and text) for sampled localities 

in the present study (sample groups as indicated). The calculated extents vary widely for most of the samples. 

Petrography‐based reaction extents are 15 and 29% higher for the more palagonitized Surtsey and Mosfell (see 

PPL photomicrograph insets) samples respectively than for the corresponding less palagonitized samples (green 

arrows). The same values determined using the SiO 2 /MgO method are 29 and 25% (violet arrows). However, those 

determined using the TiO 2 method are −5 and −15% (blue arrows). 

1985; Jakobsson and Moore, 1986]. Crovisier et al. 

[1992] found palagonitized rind thickness not to be a 

good indicator of reaction progress. Studies using 

major‐element concentration data as the basis for 

reaction extent include Staudigel and Hart [1983], 

Crovisier et al. [1992], and Stroncik and Schmincke 

[2001], in which assumed passive accumulation of 

TiO 2 , SiO 2 /MgO, or MgO, respectively, was used 

as the basis for determining elemental changes 

during palagonitization. Since the present study 

shows that TiO 2 , SiO 2 , and MgO mass change 

behaviors vary significantly among different samples 

(Table 7), palagonitization extent values based 

on these element behaviors cannot be used to 

compare palagonitization extent among samples. 

Furthermore, SiO 2 , TiO 2 and particularly MgO 

concentrations vary as a function of analysis location 

within a given palagonite rind due to the 

observed concentration gradients for these elements 

(Figures 12 and S3). 

[23] Since palagonitization extent could potentially 

depend on many factors, including initial sideromelane 

composition, palagonitization duration, 

available heat, water/rock (i.e., porosity), fluid 

chemistry, and analysis location, a petrographicbased 

approach to determining the extent of palagonitization, 

representing an average extent for 

the sample, rather than a chemically based extent 

of palagonitization seems appropriate for sample 

comparison. Here we suggest an alternative, 

petrography‐based approach, wherein palagonitization 

extent is calculated based on observed modal 

data (see section 3.1 and Table 2). Palagonitization 

extents calculated from both modal data (this study) 

and using the major‐element concentration data 

approach used by previous workers are compared in 

Figure 13; wide variations are apparent. At Surtsey 

and Mosfell, petrographically distinct (Figure 13, 

insets) samples were collected from areas that 

were less‐ and more‐palagonitized, based on field 

observations. Calculated extents using the SiO 2 / 

MgO method correlate with the petrographically 

observable difference in degree of palagonitization 

between the samples; the TiO 2 method does not. 

However, as mentioned above, the SiO 2 /MgO 

method is shown here to depend on analyzed spot 

location, particularly for MgO. The petrographic 

palagonitization extent provides a basis for comparison 

of other measured parameters in samples 

from different locations that are known to have 

palagonitized under different conditions. For 

example, in this study, palagonite water content is 


Geochemistry 

Geophysics 


observed to decrease linearly with increasing palagonitization 

extent (Figure 8); this finding is discussed 

further below. 

5.2. Porosity Control of Palagonitization 

Extent 

[24] The observed general trend of increasing 

palagonitization extent with decreasing minuscement 

porosity (i.e., decreasing original porosity; 

Figure 3) across the sample set suggests that porosity 

plays a controlling role in the extent of palagonitization. 

Walton and Schiffman [2003] made a similar 

finding for palagonitized HSDP hyaloclastites. Less 

fluid flow would be expected within low‐porosity 

samples (due to the expected relative lack of permeability), 

so we can speculate that water/rock ratio 

(w/r) decreases with decreasing porosity. Therefore 

the trend observed in Figure 3 suggests that sample 

w/r is inversely related to palagonitization extent. 

Pauly [2011] showed that pH (presumably determined 

by sample w/r and thus porosity) controls B 

isotope fractionation between palagonite and seawater. 

In that study, palagonitized sideromelane 

samples having the lowest zeolite abundance (suggesting 

low palagonitization extent) also had the 

lightest d 11 B values, as would be expected for 

samples that palagonitized in relatively low pH, 

more porous environments (presumably where little 

if any porosity reduction due to burial has occurred). 

This finding confirms that porosity is an important 

control of palagonitization extent. 

[25] It is also apparent in Figure 3 that within the 

lacustrine phreatomagmatic and subglacial volcanic 

sample groups, palagonitization extent varies 

widely among samples with similar porosity. Palagonitization 

extents of 84% and 40% were measured 

at Pahvant Butte and Black Point (lacustrine 

phreatomagmatic samples) respectively, yet these 

samples have nearly identical porosities (Table 2). 

Evidently, within these sample groups, other factor(s) 

besides porosity, presumably including available 

heat and palagonitization duration, exert control on 

palagonitization extent. 

5.3. Non‐isovolumetric Process: 

Implications for Major‐Element Changes 

[26] Previous workers [Hay and Iijima, 1968a, 

1968b; Honnorez, 1967, 1972; Eggleton and Keller, 

1982; Staudigel and Hart, 1983; Furnes, 1984; 

Bednarz and Schmincke, 1989; Zhou and Fyfe, 

1989; Jercinovic et al., 1990; Thorseth et al., 

1991; Crovisier et al., 1992; Daux et al., 1994; 

Stroncik and Schmincke, 2001; Walton et al., 2005] 

studied major‐ and/or trace element changes during 

palagonitization as a means of understanding the 

geochemical mechanism of palagonitization, using 

various methods and assumptions. Stroncik and 

Schmincke [2001] and Walton and Schiffman 

[2003] summarize these studies. This impressive 

number of studies, spanning nearly 40 years, gives 

an indication of the (still) elusive nature of elemental 

mobility during palagonitization. 

[27] In the present study, the availability of both 

major‐ and trace‐element data, obtained using in situ 

microanalysis at essentially the same spots within 

sampled palagonites, is exploited to determine 

major‐element and overall mass changes during 

palagonitization, using the isocon method [Grant, 

1986]. The observed large major‐element mass changes 

(Table 7) are well in excess of changes in solidphase 

specific gravity that can be attributed to 

hydration alone, and require extensive dissolution 

and creation of microporosity. This indicates that 

palagonitization cannot be an isovolumetric process, 

as previously suggested based on petrographic evidence 

[Walton and Schiffman, 2003; Stroncik and 

Schmincke, 2001; Jercinovic et al., 1990]. Even 

though, when observed petrographically, preserved 

primary texture and absence of secondary pores 

within palagonitized areas suggests an isovolumetric 

process, evidently significant microporosity develops 

during palagonitization, potentially in the leached 

layers imaged within gel palagonite by Drief 

and Schiffman [2004]. These workers concluded 

that the leached layers result from selective dissolution, 

and form palagonite. The finding that palagonitization 

is not an isovolumetric process calls 

into question element changes that have previously 

been calculated in weight per unit volume [Daux et al., 

1994; Crovisier et al., 1992; Hay and Iijima, 1968a], 

and suggests instead that the isocon approach of Grant 

[1986], in which mobile element concentration changes 

are determined based on immobile element concentration 

changes, is necessary to understand mass 

transfer during palagonitization. 

5.4. Major‐Element Changes: Implications 

for Seawater Cation Flux 

[28] Major‐element concentration changes determined 

from immobile‐element isocon lines in the 

present study (Table 7) confirm that palagonitization 

of sideromelane can be a significant source (or 

sink) of oceanic cations. The averaged element 

changes determined for oceanic samples in the 

present study are for the most part in reasonable 

agreement with elemental changes calculated for 


Geochemistry 

Geophysics 


palagonitization of MORB crust by Staudigel and 

Hart [1983] (Table 7). The lone exception is 

K 2 O, for which average change calculated for 

oceanic samples in the present study is −12%, 

whereas Staudigel and Hart [1983] calculated 

a value of 900%. An important observation is 

that both overall and specific major‐element mass 

changes depend strongly on original sideromelane 

composition. In particular, averaged TiO 2 and FeO 

changes range from 1.3 and −25.7% respectively 

during palagonitization of subalkaline sideromelane 

(i.e., neutral to a source of seawater cations) to 19.5 

and 6.4% respectively during palagonitization of 

alkaline sideromelane (i.e., sink of oceanic cations). 

This finding points out the need to account for variable 

mass changes due to sideromelane composition 

when performing global oceanic flux calculations. 

5.5. Compositional Control 

[29] Palagonitized rinds on the subalkaline and 

alkaline grains of Hilina Bench sample S508–7 were 

observed to have distinguishable major‐element 

compositions, even though they were palagonitized 

under identical conditions (Table 6). In addition, the 

average overall mass loss values across the entire 

sample set are very different for palagonitization 

of alkaline (26%) versus subalkaline (48%) sideromelane 

(Table 7). It is clear that palagonite 

composition depends on original sideromelane 

composition. This finding is potentially inconsistent 

with Stroncik and Schmincke [2001], who found 

palagonite composition to be mainly controlled by 

palagonite aging, beginning with the first appearance 

of palagonitized sideromelane. However, it 

should be noted that for the most part the samples 

analyzed by Stroncik and Schmincke [2001] are 

significantly older than those of the present study; 

also all samples analyzed in the present study are gel 

palagonite. 

5.6. Subalkaline Versus Alkaline 

Sideromelane Dissolution Rate 

[30] The substantial difference in average overall 

mass loss values (Table 7) indicate that subalkaline 

sideromelane dissolves much more readily than 

alkaline. This observation cannot be attributed to 

differences in palagonitization duration or conditions, 

because subalkaline versus alkaline sideromelane 

grains in Hilina Bench sample S508–7, 

which must have palagonitized for identical duration 

and conditions, lost 35.9% and 13.0% mass, 

respectively (Figure 10 and Table 7). The thicker 

versus thinner palagonitized rinds that developed 

on the subalkaline versus alkaline grains in this 

sample (Figure 6) are therefore evidence that subalkaline 

sideromelane must dissolve at a relatively 

higher rate than alkaline sideromelane during 

palagonitization. This is strong evidence that sideromelane 

dissolution rate depends on sideromelane 

composition. An implication of this finding is that 

palagonitized rind thickness gives a true indication 

of palagonitization duration, as suggested by Moore 

[1966], only if the original sideromelane compositions 

in the sample set are identical. 

5.7. Water Content 

[31] The observation that palagonite water content 

decreases linearly with palagonitization extent 

(Figure 8) holds for samples from all of the environmental 

groups, including localities known or 

inferred to have palagonitized over a relatively short 

duration (volcanic and phreatomagmatic environments) 

and localities known or inferred to have 

palagonitized over a relatively long duration (submarine 

volcaniclastic environments). In particular, 

water contents of the less‐ and more‐palagonitized 

samples from both Surtsey and Mosfell (subglacial 

volcanic group) are higher and lower, respectively 

(Figure 8, blue and green arrows). Presumably 

palagonitization duration is similar for each sample 

pair from Surtsey and Mosfell. This indicates that 

water content depends on palagonitization extent 

(or possibly temperature), for a given palagonitization 

duration (assuming water content to be stable 

post‐palagonitization). Drief and Schiffman [2004, 

Figure 2d] showed that palagonite consists of 

amorphous and smectite‐like areas. The higher 

water contents observed in relatively low palagonitization 

extent samples in the present study might 

reflect a higher abundance of amorphous (and 

presumably more hydrated) material relative to 

smectite‐like areas. Similarly, the lower water contents 

observed in relatively high palagonitization 

extent samples might reflect higher degrees of crystallinity, 

consistent with inferred recrystallization of 

palagonite to smectite (see section 5.9). 

5.8. Two Distinct Palagonitization Styles 

[32] The submarine volcaniclastic samples have 

remarkably similar low porosity values, high zeolite 

abundances, thick, smooth, gel‐like palagonite rind 

textures, and high palagonitization extents (Figure 2 

and Table 2). The submarine volcanic, phreatomagmatic 

(both marine and lacustrine) and subglacial 

volcanic samples generally are much more 

variable in terms of these same parameters (Figure 3 


Geochemistry 

Geophysics 


and Table 2); palagonitization extent within these 

groups is particularly variable. These observations 

suggest two distinguishable and environmentally 

dependent palagonitization styles. 

[33] Burial‐diagenetic palagonitization. Palagonitization 

in submarine volcaniclastic localities is 

known, or is inferred, to have occurred significantly 

later than initial eruption, fragmentation, and deposition 

of hyaloclastites [e.g., Walton and Schiffman, 

2003]. In particular, the rounded sideromelane 

grains deposited in the volcaniclastic basin at Hilina 

Bench indicate that palagonitization must have 

occurred later (i.e., after transport) than the initial 

eruption and deposition of hyaloclastite. These 

samples have been buried; burial‐diagenesis is the 

likely source of thermal energy to drive palagonitization 

at these localities. Walton and Schiffman 

[2003] attributed the high degree of palagonitization 

observed near the bottom of the core to be the 

result of burial‐diagenesis. Presumably, as samples 

at submarine volcaniclastic localities underwent 

progressive burial (initial porosity loss, decrease in 

w/r), the temperature would have been relatively 

low; for example at HSDP palagonitization occurred 

at approximately 15°C [Walton and Schiffman, 

2003]. The rate of palagonitization would be 

expected to have been relatively low, but the duration 

of palagonitization might have been relatively 

long; Walton and Schiffman [2003] estimated that 

100 mm thick palagonitized hyaloclastite rinds 

developed over a 5000‐year interval in their HSDP 

samples. These factors in combination would be 

expected to produce the observed relatively high 

palagonitiation extent, smooth, complete palagonitized 

rinds, relatively low water content, and relatively 

high REE enrichment in all of these samples. 

[34] Hydrothermal palagonitization. On the other 

hand, palagonitization of the seafloor volcanic, 

phreatomagmatic, and subglacial volcanic samples 

is known or inferred to have occurred essentially 

syn‐eruptively or relatively shortly thereafter. The 

relatively high temperatures associated with the 

volcanic activity at these localities would be 

expected to establish a potentially vigorous hydrothermal 

system, flushing heated fluid through the 

relatively porous volcanic edifice. Palagonitization 

at these environments would be expected to have 

occurred at a higher rate (presumably due to the 

active hydrothermal system) as compared to palagonitization 

in submarine environments. For example 

at the surface of Surtsey, Jakobsson and Moore 

[1986] and Jakobsson et al. [2000] mapped rapid 

palagonitization (advancement of consolidated 

tephra areas visible during annual field visits) that 

was dependent on the volcano’s hydrothermal system. 

Palagonitization duration at these localities 

would be expected to be relatively short compared to 

that expected for submarine volcaniclastic systems, 

due to dependence on the longevity of the volcanic 

heat source. Palagonitized samples from these 

environments therefore would be expected to be 

distinguishable from submarine volcaniclastic samples 

in terms of palagonitized rind texture (relatively 

thin, less well‐developed), zeolite abundance (relatively 

low). Palagonitization extents for these samples 

would be expected to be relatively low, 

resulting in higher water contents as was generally 

observed. 

[35] Of special note however are the rind textures 

(complete and fairly smooth), zeolite abundances 

(high; Table 2), and palagonitization extents (high; 

Table 2) for the Ingólfsfjall (subglacial volcanic) and 

Pahvant Butte (lacustrine phreatomagmatic) samples; 

in terms of these parameters, these samples resemble 

submarine volcaniclastic samples. The observed 

greater palagonitization extent observed for these 

samples potentially could be due to exceptionally 

long palagonitization duration or exceptionally high 

palagonitization temperatures. Another possibility is 

a second stage of palagonitization that could occur in 

seafloor volcanic, marine phreatomagmatic, and 

subglacial volcanic environments due to subsequent 

burial‐diagenesis of the volcanic edifice. Jakobsson 

and Gudmundsson [2008] found that hydrothermal 

activity caused tephra palagonitization and consolidation 

of the edifices during the first years after the 

eruptions of Surtsey (1963–1967) and Gjálp (1996), 

but that alteration and consolidation of the outer 

slopes likely occurs over thousands of years and is 

mainly a diagenetic process. In summary, burialdiagenesis 

and hydrothermal palagonitization styles 

are distinguishable based on palagonitized rind 

petrographical and in situ analytical observations, 

however care must be taken interpreting the palagonitization 

of samples from seafloor volcanic, 

phreatomagmatic, and subglacial volcanic environments. 

Unambiguous distinction of palagonitization 

style requires additional information regarding 

the potential occurrence of post‐eruptive burialdiagenesis 

in these environments. 

5.9. Formation and Evolution 

of Alteration Layers 

[36] Palagonitization arguably must begin at the 

outer margin of fresh sideromelane shards, where 

initially sideromelane is in contact with fluid and 

begins to dissolve, and then proceed inward with 


Geochemistry 

Geophysics 


Figure 14. Schematic model depicting evolution and migration of sideromelane alteration layers. Arrows indicate 

evolution direction of the gel‐palagonite and phyllosilicate alteration layers with time. The contact between the phyllosilicate 

and gel‐palagonite layers likely is gradational (purple area). The observed SiO 2 ,Al 2 O 3 , and CaO (mobile) 

and Zr, Nb, Sc, La, Nd (immobile) concentration gradients are consistent with progressive sideromelane dissolution 

with time, however the observed MgO (mobile element) concentration gradient is not. The MgO concentration gradient 

is consistent with the syn‐ and/or post palagonitization development of a phyllosilicate (e.g., smectite) layer, 

forming as a consequence of the progressive evolution of the gel‐palagonite layer to equilibrium with the fluid. 

The inferred degree of development of smectite is highest at the fluid boundary. 

time; Walton and Schiffman [2003] showed this to 

be true for samples of palagonitized hyaloclastite 

from HSDP. The observed palagonite REE concentration 

gradients in the present study (Figures 9 

and S1) suggest therefore that REE concentration 

changes with time. This finding contrasts with 

Daux et al. [1994], who found that changes in 

palagonite REE concentration are independent of 

reaction progress. Because we have found that the 

REEs are immobile during palagonitization, mobileelement 

mass loss must be greatest at the outer 

palagonitized margin, and decrease inward toward 

sideromelane. Furthermore, dissolution must continue 

behind the palagonitization front in order to account 

for the observed REE concentration gradients. 

[37] Because MgO was found to be highly mobile 

(Table 7), the observed MgO concentration gradients 

(decreasing through palagonite toward sideromelane; 

Figures 12 and S3) are inconsistent 

(opposite direction of) with gradients that would be 

expected due to cumulative dissolution with time ‐ 

the MgO concentration gradients do not record 

sideromelane dissolution. Rather, they are consistent 

with syn‐ and/or post‐palagonitization conversion 

of gel‐palagonite to phyllosilicate (e.g., 

smectite), beginning at the outer palagonitized 

margin and proceeding inward toward sideromelane 

with time. Valle et al. [2010] documented the formation 

and inward‐thickening with time of a phyllosilicate 

layer at the outer margin of gel‐palagonite 

layers that developed on synthetic borosilicate 

(nuclear waste containment) glass, and these workers 

suggest that similar behavior during alteration of 

natural sideromelane would be expected. A schematic 

model depicting the evolution and migration 

of alteration layers in this framework is given in 

Figure 14. Valle et al. [2010] concluded that the gel 

(leached) layer developed on borosilicate glass surfaces 

is not at equilibrium with the solution, 

implying that this layer evolves with time. Honnorez 

[1967, 1972, 1978] and Stroncik and Schmincke 

[2001] concluded that the ultimate fate of gelpalagonite 

is conversion to smectite, and Drief and 

Schiffman [2004] documented both amorphous and 

crystalline areas within the outer palagonitized 

margins of natural sideromelane samples. The MgO 

concentration gradients within palagonites observed 

in the present study are consistent with these previous 

studies, and likely represent the continuing 


Geochemistry 

Geophysics 


evolution (conversion) of the gel‐palagonite layer 

toward a phyllosilicate layer, which would be 

expected to be approaching equilibrium with the 

fluid. 

6. Conclusions 

[38] 1. Palagonite water content decreases with 

increasing palagonitization extent, likely due to 

recrystallization of palagonite into smectite. 

[39] 2. Palagonitization extent varies inversely with 

original sample porosity, confirming that porosity 

(i.e., associated permeability) is an important controlling 

factor of palagonitization; however, other 

factors are also important. 

[40] 3. Zr, Nb, Sc, and REE concentrations in 

palagonite and sideromelane plot along isocon 

lines, indicating that these elements are immobile 

during palagonitization. 

[41] 4. Large major‐element mass changes indicate 

that palagonitization entails extensive dissolution 

and creation of microporosity, and therefore cannot 

be considered an isovolumetric process, as has 

been assumed previously by many workers. 

[42] 5. Major‐element mass changes due to palagonitization 

for SiO 2 ,Al 2 O 3 , and in particular TiO 2 

and FeO depend on the subalkaline versus alkaline 

character of the sideromelane. 

[43] 6. Original sideromelane composition is an 

important control of palagonite composition. 

Therefore it is necessary to take original sideromelane 

composition (and resulting expected 

mass loss for specific elements) into account when 

calculating estimates of global seawater cation flux. 

[44] 7. Subalkaline sideromelane appears to dissolve 

at a much higher rate than alkaline sideromelane 

during palagonitization. 

[45] 8. Two distinct palagonitization styles, burialdiagenetic 

and hydrothermal, are recognized. 

[46] 9. Observed palagonite REE concentration 

gradients indicate that sideromelane dissolution 

continues in the zone behind the advancing palagonitization 

front. 

[47] 10. Observed palagonite MgO concentration 

gradients appear to reflect the early stages of synand/or 

post‐palagonitization conversion of the gelpalagonite 

layer to a phyllosilicate (e.g., smectite) 

layer, consistent with the evolution of sideromelane 

alteration layers toward equilibrium with the surrounding 

fluid. 

Acknowledgments 

[48] Michelle Coombs, Mike Garcia, and Don Thomas graciously 

provided some of the samples used in this study, as well 

as thoughtful discussions. José Honnorez provided the Palagonia 

sample. Roy Bailey and James White provided very helpful 

field guidance at Black Point and Pahvant Butte, respectively. 

Svienn Jakobsson and The Surtsey Committee granted access 

to Surtsey, and Hjalti Franzson supported and participated in 

the Iceland fieldwork. Tony Walton and Jim Moore provided 

very useful discussion support. Jake Lowenstern provided lab 

access and guidance for the reflected‐light FTIR measurements. 

Alicé Davis provided expert assistance with petrographic imaging; 

Norm Winter and Greg Baxter provided thin section sample 

preparation. Sarah Roeske and Brian Joy were of tremendous 

help during the extensive electron microprobe sessions, and 

Michelle Gras was equally helpful in the laser‐ablation ICPMS 

lab. This work was supported by NSF‐EAR grant 0125666. The 

authors thank José Honnorez and Tony Walton for very helpful 

reviews of the initial manuscript. 

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