# SUBAQUEOUS VOLCANISM, FROM ANCIENT SUCCESSIONS TO MODERN VOLCANOES AND MODELLING

EDITED BY : Carles Soriano, Nancy Riggs and Guido Giordano PUBLISHED IN : Frontiers in Earth Science

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ISSN 1664-8714 ISBN 978-2-88963-699-0 DOI 10.3389/978-2-88963-699-0

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# SUBAQUEOUS VOLCANISM, FROM ANCIENT SUCCESSIONS TO MODERN VOLCANOES AND MODELLING

Topic Editors:

Carles Soriano, Instituto de Ciencias de la Tierra Jaume Almera (ICTJA), Spain Nancy Riggs, Northern Arizona University, United States Guido Giordano, Roma Tre University, Italy

Citation: Soriano, C., Riggs, N., Giordano, G., eds. (2020). Subaqueous Volcanism, From Ancient Successions to Modern Volcanoes and Modelling. Lausanne: Frontiers Media SA. doi: 10.3389/978-2-88963-699-0

# Table of Contents

*04 Editorial: Subaqueous Volcanism, From Ancient Successions to Modern Volcanoes and Modelling*

Carles Soriano, Nancy Riggs and Guido Giordano

William W. Chadwick Jr., Susan G. Merle, Edward T. Baker, Sharon L. Walker, Joseph A. Resing, David A. Butterfield, Melissa O. Anderson, Tamara Baumberger and Andra M. Bobbitt

David A. Clague, Jennifer B. Paduan, David W. Caress, Craig L. Moyer, Brian T. Glazer and Dana R. Yoerger

Danilo Cavallaro and Mauro Coltelli

# Editorial: Subaqueous Volcanism, From Ancient Successions to Modern Volcanoes and Modelling

Carles Soriano1 \*, Nancy Riggs 2 \* and Guido Giordano3 \*

*1 Instituto de Ciencias de la Tierra Jaume Almera, CSIC, Barcelona, Spain, 2 Merriam-Powell Center for Environmental Research, Northern Arizona University, Flagstaff, AZ, United States, 3 Department of Science, Roma Tre University, Rome, Italy*

Keywords: processes, products, underwater eruptions, underwater vehicles, bathymetry

#### **Editorial on the Research Topic**

#### **Subaqueous Volcanism, From Ancient Successions to Modern Volcanoes and Modelling**

Understanding the processes and products of subaqueous volcanism has been traditionally based on the study of ancient successions and deposits; only a few studies are based on deposits from recent subaqueous eruptions and fewer yet on the direct observation of underwater eruptions. This situation has rapidly changed during last decades of the twentieth century and first decades of the twenty-first century, during which scientific and technological development has allowed an increasing exploration of the sea floor. Understanding of subaqueous volcanism has benefited from such exploration and from an increasing number of studies on modern volcanoes under the sea. As a result, the processes and products of subaqueous eruptions are now much better understood than ever before. Nevertheless, the study of ancient volcanic successions around the world still is a major source of models for subaqueous eruptions and, together with experimental and numerical modeling, is a major contribution to our understanding.

The present Research Topic showcases current trends in research on subaqueous volcanism, with a dominance of oceanographic surveys aimed toward the study of modern volcanoes and their deposits by using a number of techniques such as different types of underwater vehicles (AUVs, ROVs) and of high-resolution sonar bathymetries.

Cas and Simmons is a review with a summary of the physical properties of water and on how these may affect magma properties and eruptive processes like magma decompression, volatile exsolution and, ultimately, volcanic explosiveness. This contribution discusses a major issue in subaqueous volcanism that has driven a long-standing controversy: the hydrostatic pressure exerted by water column and its capability to inhibit explosive eruptions. It also addresses a process that has been increasingly investigated in both subaerial and subaqueous volcanic studies, as is the non-explosive production of ash by pumice-to-pumice abrasion.

The Havre volcano in the Kermadec arc erupted in 2012 and is likely one of the best-documented recent deep-water eruptions on Earth. Two contributions in the Research Topic are devoted to the study of this eruption of silicic magma. Ikegami et al. focuses on the chronostratigraphic relations of lavas and domes to reconstruct the chronologic history of the Havre eruption based on vent distribution and on lava and dome stratigraphy. They also provide a critical comparison of the Havre lavas and domes with subaerial counterparts, finding close similarities between them.

Murch et al. analyzed the Ash with Lapilli unit from Havre volcano by means of ROV sampling and detailed imagery. They use their results to highlight the important role that small-volume ash beds play in understanding subaqueous volcanic activity as a whole.

Edited and reviewed by: *Valerio Acocella, Roma Tre University, Italy*

#### \*Correspondence:

*Carles Soriano<EMAIL_ADDRESS>Nancy Riggs<EMAIL_ADDRESS>Guido Giordano<EMAIL_ADDRESS>
#### Specialty section:

*This article was submitted to Volcanology, a section of the journal Frontiers in Earth Science*

Received: *15 January 2020* Accepted: *04 February 2020* Published: *21 February 2020*

#### Citation:

*Soriano C, Riggs N and Giordano G (2020) Editorial: Subaqueous Volcanism, From Ancient Successions to Modern Volcanoes and Modelling. Front. Earth Sci. 8:42. doi: 10.3389/feart.2020.00042*

**4**

Clague et al. report on the construction and destruction processes in a volcanic seamount to the southeast of Hawaii Island. Construction events were followed by vertical collapses due to lateral migration of basaltic magma to feed flank rift eruptions, yielding morphology of nested calderas, pit craters, and resurgent caldera blocks. Down-talus emplacement of basaltic pillows during eruption and syn-eruptive landslides on volcano flanks modified volcanic edifices.

Chadwick et al. present documentation of the first known historical eruption in the Mariana spreading center in December 2015, with a depth of >4,000 m below sea level. Multiple surveys, both from ships and using submersible ROVs, let the authors pinpoint the timing of eruption within a 3-year period. Although the Mariana system has a slow spreading rate, the study concludes that an active magmatic system is clearly present.

Cavallaro and Coltelli investigates the morphology of the Graham Volcanic Field (Sicily Channel, Italy) based on High-Resolution Seafloor Mapping and ROV Images. Their results show the tectonic control on the N-S and NW-SE alignment of monogenetic cones, which include the remnant of the 1,831 submarine-to-emergent eruption that formed the now totally dismantled Ferdinandea Island.

Rizzo et al. analyzed the geochemistry of CO2-rich gases associated with the Kolumbo submarine volcano, which lies next to Santorini (Hellenic Volcanic Arc, Greece). Their results show that magmatic gases equilibrate within the Kolumbo hydrothermal system at about 270◦C and at a depth of ∼1 km b.s.l.

Goto and Tomiya's contribution discusses a Quaternary subaerial cryptodome, highlighting commonalities and differences with respect to domes that grow in underwater settings.

These contributions have improved our empirical and theoretical understanding of the processes and products of subaqueous volcanism. As scientific and technological development proceeds, more studies of modern sea-floor volcanoes will provide empirical data on which refinement of theoretical models will be based. This assures that our understanding of subaqueous volcanism will continue to improve.

### AUTHOR CONTRIBUTIONS

All authors listed have made a substantial, direct and intellectual contribution to the work, and approved it for publication.

**Conflict of Interest:** The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Copyright © 2020 Soriano, Riggs and Giordano. This is an open-access article distributed under the terms of the Creative Commons Attribution License (CC BY). The use, distribution or reproduction in other forums is permitted, provided the original author(s) and the copyright owner(s) are credited and that the original publication in this journal is cited, in accordance with accepted academic practice. No use, distribution or reproduction is permitted which does not comply with these terms.

## The Eruption of Submarine Rhyolite Lavas and Domes in the Deep Ocean – Havre 2012, Kermadec Arc

Fumihiko Ikegami1 , Jocelyn McPhie1 \*, Rebecca Carey1 , Rhiannan Mundana1 , Adam Soule2 and Martin Jutzeler1

1 School of Natural Sciences, University of Tasmania, Hobart, TAS, Australia, 2 Woods Hole Oceanographic Institution, Woods Hole, MA, United States

#### Edited by:

Carles Soriano, Instituto de Ciencias de la Tierra Jaume Almera (ICTJA), Spain

#### Reviewed by:

Jonathan Fink, Portland State University, United States Ilya Bindeman, University of Oregon, United States

> \*Correspondence: Jocelyn McPhie<EMAIL_ADDRESS>
#### Specialty section:

This article was submitted to Volcanology, a section of the journal Frontiers in Earth Science

Received: 29 June 2018 Accepted: 14 September 2018 Published: 17 October 2018

#### Citation:

Ikegami F, McPhie J, Carey R, Mundana R, Soule A and Jutzeler M (2018) The Eruption of Submarine Rhyolite Lavas and Domes in the Deep Ocean – Havre 2012, Kermadec Arc. Front. Earth Sci. 6:147. doi: 10.3389/feart.2018.00147 Silicic effusive eruptions in deep submarine environments have not yet been directly observed and very few modern submarine silicic lavas and domes have been described. The eruption of Havre caldera volcano in the Kermadec arc in 2012 provided an outstanding database for research on deep submarine silicic effusive eruptions because it produced 15 rhyolite (70–72 wt.% SiO2) lavas and domes with a total volume of ∼0.21 km3 from 14 separate seafloor vents. Moreover, in 2015, the seafloor products were observed, mapped and sampled in exceptional detail (1-m resolution) using AUV Sentry and ROV Jason2 deployed from R/V Roger Revelle. Vent positions are strongly aligned, defining NW-SE and E-W trends along the southwestern and southern Havre caldera margin, respectively. The alignment of the vents suggests magma ascent along dykes which probably occupy faults related to the caldera margin. Four vents part way up the steeply sloping southwestern caldera wall at 1,200–1,300 m below sea level (bsl) and one on the caldera rim (1,060 m bsl) produced elongate lavas. On the steep caldera wall, the lavas consist of narrow tongues that have triangular cross-section shapes. Two of the narrow-tongue segments are connected to wide lobes on the flat caldera floor at ∼1,500 m bsl. The lavas are characterized by arcuate surface ridges oriented perpendicular to the propagation direction. Eight domes were erupted onto relatively flat sea floor from vents at ∼1,000 m bsl along the southern and southwestern caldera rim. They are characterized by steep margins and gently convex-up upper surfaces. With one exception, the domes have narrow spines and deep clefts above the inferred vent positions. One dome has a relatively smooth upper surface. The lavas and domes all consist of combinations of coherent rhyolite and monomictic rhyolite breccia. Despite eruption from deep-water vents (most >900 m bsl), the Havre 2012 rhyolite lavas and domes are very similar to subaerial rhyolite lavas and domes in terms of dimensions, volumes, aspect ratio, textures and morphology. They show that lava morphology was strongly controlled by the pre-existing seafloor topography: domes and wide lobes formed where the rhyolite was emplaced onto flat sea floor, whereas narrow tongues formed where the rhyolite was emplaced on the steep slopes of the caldera wall.

Keywords: lava, dome, submarine effusive eruption, rhyolite, Havre

### INTRODUCTION

feart-06-00147 October 15, 2018 Time: 19:28 # 2

Silicic effusive eruptions in deep submarine environments have not yet been directly observed, partly because the deep sea floor is a difficult environment to explore and partly because such eruptions are apparently rare. The majority of our knowledge comes from the study of silicic lavas and domes in ancient submarine successions now exposed on land (e.g., Cas, 1978; Yamagishi and Dimroth, 1985; Kano et al., 1991). However, such studies have yielded only limited information on the lava morphology and dimensions; also, the duration of lava emplacement, the details of the original setting and the water depth at the vent are typically poorly constrained.

Recently, autonomous underwater vehicles (AUVs) and remotely operated vehicles (ROVs) have allowed high-resolution observation and strategic sampling of silicic lavas on the modern sea floor, adding significantly to data on their characteristics. For example, Allen et al. (2010) utilized ROV to document the morphological and textural differences among rhyolite lava domes at Sumisu-I knoll, Izu-Bonin arc. At Myojin-knoll caldera, 70 km farther north, Honsho et al. (2016) produced a highresolution AUV map of an intracaldera rhyolite dome. A series of studies in the Manus Basin, Papua New Guinea, including highresolution bathymetry, seafloor observation, and sampling by underwater vehicles as well as drilling (Ocean Drilling Program Leg 193), provided a detailed description of a deep submarine dacite dome and lava complex on the modern sea floor (Bartetzko et al., 2003; Paulick et al., 2004; Binns et al., 2007; Thal et al., 2014). Embley and Rubin (2018) described dacitic lavas and domes in the northeastern Lau Basin that had been explored using multibeam sonar, camera tow and dredging.

This paper extends our understanding of deep submarine silicic lavas by adding examples produced during the 2012 eruption at Havre caldera, the first of such eruptions to be witnessed (Carey et al., 2018). The eruption produced 15 separate rhyolite lavas and domes from 14 vents between ∼880 m and 1,280 m bsl. In some cases, rhyolite was erupted onto flat sea floor and in others, it traversed steep slopes. The lavas and domes have been mapped using exceptionally high-resolution bathymetry, observed by video and still photography, and sampled. We present detailed information on dimensions, volumes, aspect ratios, textures, surface features, and morphology, and explore the influence of vent depth and substrate slope on lava morphology. We also compare these submarine rhyolite lavas and domes with subaerial counterparts, and demonstrate that despite eruption in more than 900 m of seawater, the Havre rhyolite lavas and domes are very similar to subaerial rhyolite lavas and domes.

#### HAVRE 2012 ERUPTION AND MESH CRUISE IN 2015

The eruption of Havre caldera volcano (31◦ 6.50 S, 179◦ 2.450 W) in the Kermadec arc in 2012 (**Figure 1A**) was first detected by the appearance of extensive floating pumice rafts on 18th July, 2012 (Global Volcanism Program, 2012; Carey et al., 2014; Jutzeler et al., 2014; Carey et al., 2018). Carey et al. (2018) estimated the bulk volume of pumice in the rafts to be ∼1.2 km3 . On the 15th October, 90 days after the production of the pumice rafts, R/V Tangaroa of NIWA (National Institute of Water and Atmospheric Research) visited the area and acquired bathymetry at Havre. In 2015, the MESH (Mapping, Exploration, & Sampling at Havre) voyage was conducted in order to perform more detailed investigation of the 2012 eruption products on the sea floor. As well as rhyolite lavas and domes, the 2012 eruption generated a layer of giant pumice (GP) clasts (GP deposit; ∼0.1 km3 bulk) and small-volume pyroclastic deposits (ash-lapilli-block deposit and ash-lapilli deposit; <0.07 km3 bulk) on the sea floor (Carey et al., 2018). Data were collected using two unmanned vehicles, Jason2 ROV and Sentry AUV from the US National Deep Submersible Facility. The AUV Sentry created a high-resolution bathymetry map covering the Havre caldera and adjacent areas. The ROV Jason2 operated in tandem, traversing the sea floor and sampling.

### Bathymetry Acquisition and Comparisons

The MESH cruise acquired two bathymetry datasets with different platforms. The EM122 multi-beam echo sounder (12 kHz) on R/V Roger Revelle provided data resolution to create a 25-m gridded Digital Elevation Model (DEM) of the Havre caldera and surrounding areas. High-resolution bathymetry data (1-m gridded DEM) were collected with a Reson SeaBat 7125 multi-beam echo sounder (400 kHz) on the AUV Sentry. The Sentry bathymetry data cover 56.5 km2 (total 465 km track; **Supplementary Figure S1**).

Two other bathymetric DEMs acquired in the past are also used for comparison. The earlier data were acquired in 2002 (Wright et al., 2006) by the R/V Tangaroa which was equipped with a Simrad EM300 multi-beam echo sounder (30 kHz). The later data were acquired by the same R/V Tangaroa in 2012, 90 days after the pumice rafts appeared in the satellite imagery (Carey et al., 2014), using a Kongsberg EM302 multi-beam echo sounder (30 kHz) which allows the creation of an approximately 24-m gridded DEM. The topographic differences between the DEMs were calculated by Raster Calculator of QGIS 2.14 after datum and projection of the grids were adjusted to Zone 1S of Universal Transverse Mercator projection. The DEMs are referenced to best fit the AUV bathymetry data.

#### Visual Observations and Sampling by ROV

The ROV Jason2 conducted 12 dives (J2-802–J2-813) and traversed 50 km of sea floor over 238.5 h. The traverses were recorded by three video cameras and two still cameras. The dives covered the floor, walls, and rims of the Havre caldera (**Supplementary Figure S2**). Dives 802–808, 811, and 813 were dedicated to the investigation of the 2012 rhyolite lavas and domes, whereas the other dives observed the eastern, northern, and western part of the caldera. Eighty-nine lava and dome samples were collected during the ROV Jason2 traverses (**Supplementary Figures S3–S5**). A subset of 41 lava and dome samples was examined by both an optical

are: Kermadec trench, Kermadec ridge (KR), Havre Trough (HT), and Colville Ridge (CR); (Wright et al., 2006). (B) Bathymetry map of Havre caldera in 2012 after the eruption (Carey et al., 2014). Approximate extent of the 2012 rhyolite lavas and domes shaded purple based on MESH 2015 data.

microscope and scanning electron microscope (SEM). Wholerock X-ray Fluorescence analyses of samples from each of the lavas and domes have been published in Carey et al. (2018).

### SETTING OF HAVRE 2012 LAVAS AND DOMES IN HAVRE CALDERA

Havre is a 5-km-wide caldera (Wright et al., 2006); the caldera rim is situated at ∼1,000 m bsl and the floor is at ∼1,500 m bsl (**Figure 1B**). Very little is known about the formation of Havre caldera. Wright et al. (2006)recovered samples of basalt, andesite, and dacite lavas and intrusions thought to form the pre-caldera sequence.

Comparison of the 2002 and 2012/2015 bathymetry shows that the 2012 eruption changed the southern and southwestern caldera margin dramatically, with the appearance of 15 new lavas and domes (**Figure 2** and **Supplementary Figure S6**). Syneruptive mass-wasting events also modified the western and southwestern caldera margin and the caldera floor. No significant changes have been recognized elsewhere on the edifice. The lavas and domes have been labeled A to P from west to east. Note that J is also a silicic dome but it was already present in 2002. The 2012 lavas and domes are rhyolitic, ranging between 70.7 and 72.5 wt.% SiO2 (Carey et al., 2018). Vents for the lavas and domes are located around the southwestern and southern margin of the caldera (**Figure 2**). The vents extend along 4.2 km of the caldera perimeter, the total length of which is 15 km. Five of the vents (A to E) are part way up the steep southwestern caldera wall and the rest (F to I, K to P) are on the southwestern and southern rim of the caldera.

### PETROGRAPHY OF THE 2012 RHYOLITE LAVAS AND DOMES

The typical texture is porphyritic (<5 modal% phenocrysts, 100– 500 µm) (**Figure 3**) and microlite-rich (>60 modal%) in glassy groundmass (**Figure 4**). The phenocryst population comprises plagioclase, pyroxene, and Fe-Ti oxides, and a minor amount of quartz and apatite. The microlites are mainly plagioclase (>70 modal%); pyroxene (∼20 modal%) and Fe-Ti oxides (∼10 modal%) make up the rest. The vesicularity varies from <5 modal% to >50 modal%. The most vesicular samples (>50 modal%; e.g., L and the carapace of G) have a pumiceous texture and are exceptionally poor (<5 modal%) in microlites although their phenocryst populations are the same as in other samples

collected. Extremely elongate vesicles and bands characterized by different microlite contents are common in the carapace of G (**Figures 3c**, **4d**); L samples are microvesicular.

Cristobalite is common in the microlite-rich samples. Cristobalite is a high-temperature devitrification or vapor-phase crystallization product (Baxter et al., 1999; Horwell et al., 2013). The size (typically ∼50 µm) and amount of cristobalite are highly variable in the Havre rhyolite samples although the abundance is commonly higher in microlite-rich samples. Cristobalite occurs in vesicles or groundmass (**Figure 4b**), and in extreme cases, cristobalite has completely replaced the groundmass (**Figure 4c**). The Havre cristobalite is identical to cristobalite found in subaerial silicic lava domes (e.g., Horwell et al., 2013; Schipper et al., 2015).

#### MORPHOLOGY AND DIMENSIONS OF THE HAVRE 2012 RHYOLITE LAVAS AND DOMES

The Havre 2012 rhyolites are divided into lavas (A, B, C, E, F) and domes (D, G to I, K to P) (**Figure 5**) based on overall morphology. All the lavas and domes consist of combinations of coherent rhyolite and monomictic rhyolite breccia. Monomictic

FIGURE 3 | Photomicrographs of Havre 2012 rhyolite lava and dome samples. The yellow bars are 1 mm long. (a) XPL image of HVR141 from C, representing typical low-vesicularity texture (<15 modal% vesicles) and glassy groundmass. Plagioclase and pyroxene phenocrysts are <1 mm and sparse (<10 modal%). (b) PPL image of HVR144 from C, showing abundant small vesicles. The groundmass contains abundant microlites. (c) XPL image of HVR099 from L, showing dark gray, black, and pale gray groundmass bands defined mainly by differences in microlite abundance. (d) XPL image of HVR090 from I, showing heterogeneous vesicularity and irregular-shaped vesicles forming a band.

rhyolite breccia comprises coarse, angular fragments of dense to pumiceous, massive or flow-banded rhyolite; the fragments are prismatic or polyhedral and bounded by curviplanar surfaces.

The lavas are narrow, elongate tongues of rhyolite that were erupted from the vents on the southwestern caldera wall and rim, and descended 300–600 m to the caldera floor (**Figure 5a**). Lavas A and C extended farther across the caldera floor (A, ∼450 m; C, ∼650 m) and also spread laterally. The lavas have relatively high aspect ratios (lateral extent:thickness; Walker, 1973; 5.5–47.7; **Supplementary Table S1**). The domes on the other hand were erupted from vents on the southwestern and southern caldera rim (**Figures 5b,c**), except for D which is on the southwestern caldera wall. The domes are round or oval in plan view although G, H, and I have been truncated along their northeastern edges by a mass-wasting scarp. The domes have relatively low aspect ratios (lateral extent:thickness; 2.0–3.8; **Supplementary Table S1**).

The combination of bathymetric differences (2002 versus 2015; **Supplementary Figure S7**) as well as the precisely mapped extents of the Havre 2012 lavas and domes, have been used to calculate their median thicknesses and volumes (**Supplementary Table S1**). Near-vent thicknesses of the lavas (A, B, C, E, F) range from 50 to 100 m, and they maintain near-uniform thickness with distance from the source down the steep caldera wall (**Figure 6a** and **Supplementary Figure S8**). Most of the domes (G–I, K–N) have maximum thicknesses between 50 and 110 m (**Figure 6b**). The exception is the largest dome pair, OP, which has a maximum thickness of 210 m. The total volume of all lavas and domes is 0.21 km3 . The largest component (>50%; 0.11 km3 ) of this volume is contained by the dome pair, OP, on the southern caldera margin (**Figure 2**). The aggregate volume erupted from vents on the southwestern caldera wall (A–F) is also significant (0.08 km3 ). In comparison, the domes on the southwestern and southern caldera rim (G–I, K–N) are very small (aggregate volume 0.02 km3 ). The smallest "dome" is L which comprises the top of a rhyolite body still largely contained in its vent and too small to measure.

#### Lavas

All the lavas traversed the ∼40◦ slope of the caldera wall (**Figure 7a**). On this steep slope, the lavas are narrow tongues, more-or-less triangular in cross-section (∼200 m across and ∼50 m high; **Figure 7b**). The surfaces of the narrow tongues are almost entirely monomictic rhyolite breccia; sparse exposures of coherent rhyolite occur exclusively at the top of the narrow

tongues. Along much of the narrow-tongue segment of A, there is a shallow depression (∼15 m across and <10 m deep; **Figure 7b**). Lavas A and C also have wide-lobe segments emplaced on the flat caldera floor (**Figures 2**, **7a,c**). The wide-lobe segments have similar rough surfaces to the narrow-tongue segments, composed of monomictic rhyolite breccia; coarsely vesicular coherent rhyolite is locally exposed beneath the breccia.

Arcuate surface ridges are present on the wide-lobe segments of A and C, and on parts of the narrow-tongue segments of A, B, C, and F (**Supplementary Figure S9**), though the best example of this feature is on A (**Figure 8**). The ridges are oriented perpendicular to the long axes of the lavas and curve away from the leading edge of the lavas. On A, the ridges are sub-parallel and have wavelengths of 10– 30 m and amplitudes of 1–2 m (**Figures 8d,e**). There is one small domain in which the pattern of ridges is slightly offset (**Figure 8b**). The pattern of ridges on the wide-lobe segment of C is complex and includes several domains where the ridges have different orientations (**Figure 2a** and **Supplementary Figure S9**). Also, the distal margin of C is indistinct where it meets variably deformed caldera-floor sediment. The emplacement of the wide-lobe segment of C and its impact on the caldera-floor sediment are the subject of a separate study.

Breakout lobes are present on the wide-lobe segment of A ("x" on **Figures 5a**, **9a,b**) and the narrow-tongue segment of E ("y" on **Figures 5a**, **9c,d**). On the western margin of A close to the start of the wide-lobe segment, there is a stack of three breakout lobes. The largest is at the base and is 150 m long and ∼15 m high; two smaller breakout lobes occur immediately above. The largest breakout lobe appears to connect directly with the narrow-tongue segment of A on the steep caldera wall. The breakout lobe on E is 50 m long, 60 m across, and 15 m high. It emerges from the crest of the narrow tongue about two-thirds of the way down the caldera wall.

#### Domes

All the domes have very rough surfaces composed of jagged and spiny coherent rhyolite (**Figures 10a,b**), surrounded by aprons of coarse breccia. The spines rise almost vertically and have widths of a few m and heights up to 20 m; their surfaces are smooth and planar or gently curved (**Figure 10c**), or very rough and irregular. In some cases, spines form a ridge 10–100 m across and ∼40 m high at the center of the dome (**Figures 10a,b**). The ridges are

FIGURE 5 | Enlarged relief maps (a–c) based on AUV bathymetry map (2015) of the Havre 2012 lavas and domes. The extent of each map is shown in the small inset map. The capital letters identify the lavas and domes following Figure 2b. The bright lines are the profiles given in Figure 6. Lower case letters "x" and 'y" indicate the areas shown in detail in Figure 9.

locations are shown on Figure 5. The caldera rim is at the left and the caldera wall and floor are at the right. The horizontal and vertical scales are in metres. The vertical axis is exaggerated 200% and adjusted to the surrounding topography. The cross-sections are aligned so all the vents are at zero. (a) The lavas that descended the caldera wall (A, B, C, E, F). (b) The domes on the caldera rim (G–I, K–OP). D and L are excluded from the comparison as they are too small. Domes O and P are shown together as the basal topography of dome P is not known.

aligned either parallel to the caldera rim (K and N) or radially (H, I, and M). On K, the central ridge is cut by a narrow, deep (∼30 m) cleft (**Figure 10b**) parallel to fissures that extend between K and M (**Figure 10a**).

The large dome pair OP has a more complicated morphology due to the two extrusions apparently sharing one vent. Dome O is slightly elongate in a NE-SW direction (**Figure 5b**). The top surface slopes gently westward and consists of numerous spines

vertical exaggeration.

and ridges. The ridges define weak horseshoe-shaped alignments opening to the west and the highest point of the dome (635 m bsl) is located on its eastern margin. The top is surrounded by steep aprons of breccia (∼40o for 200 m of height). Dome P sits on the top of O and its highest point is 50 m higher than the top of O. P is elongate, being slightly longer in the northeasterly direction (∼450 m) than the northwesterly direction (∼300 m). The southwestern margin of P overlaps that of O, making their boundaries hard to distinguish. The surface of P slopes gently and is composed of decimeter-sized rhyolite fragments which

produce a smooth morphology on the AUV bathymetric map (**Figure 5b**).

Dome G is the only dome with arcuate surface ridges, similar to the arcuate surface ridges on lava A in terms of wavelength (10– 20 m) and amplitude (∼2 m). The carapace of G was sampled at the top of a mass-wasting scarp that faces the caldera, and

about 120 m Southwest of the scarp (**Figure 11**). The carapace is vesicular and glassy (**Figure 4d**). The interior of G was observed and sampled on the mass-wasting scarp. The vesicular coherent rhyolite extends ∼2 m into the interior. Dome F also has a vesicular carapace. The inner part of G consists of massive coherent rhyolite that grades deeper down into jointed coherent rhyolite.

Most of the surface of N is rough and jagged, similar to the other domes, but part of its northern edge is smoother and shows subtle tonal differences (**Figure 12**). This northern area was uplifted ∼100 m between 2002 and 2012, and both morphologically and visually, the surface resembles the surface of the nearby caldera rim, rather than the rough rhyolite. One explanation is that N is partially intrusive, a portion of the northern side of the dome having stalled and intruded instead of emerging at the sea-floor. Intrusion resulted in subtle local uplift of the caldera rim that can be detected in the detailed bathymetry.

### STRATIGRAPHIC RELATIONSHIPS OF HAVRE 2012 LAVAS AND DOMES

The order of emplacement of the 15 rhyolite lavas and domes has been deciphered using a combination of stratigraphic relationships with other 2012 seafloor products and contact relationships between adjacent lavas or domes. Among the 2012 seafloor products, the deposit of GP clasts is most useful because it was formed by settling from suspension of temporarily buoyant pumice clasts (Carey et al., 2018) and serves as a stratigraphic marker. This deposit is distinguishable in the high-resolution bathymetry map and was sampled in many places. One limitation of using the GP deposit as a marker is that its dispersal area is strongly elongate in a northwesterly direction (Carey et al., 2018) and some domes (G, H, I) on the caldera rim are outside the dispersal area. A second limitation is that the GP clasts can be found only on flat sea floor so relationships could not be identified for B, D, and E, all of which are limited to the steep caldera wall. For F, the top is on the caldera rim but outside the GP dispersal area, and the rest of F is on the steep caldera wall. Nevertheless, among the 15 lavas and domes, only A and C are overlain by numerous clasts of GP whereas K, L, M, N, and OP (all within the dispersal area and on flat sea floor) are not. Considering A in more detail, a small breakout lobe on its northwestern margin is not covered by GP clasts (**Figures 9a,b** and **Supplementary Figure S10**), implying that the breakout occurred after the GP deposit. Therefore, most of A and all of C pre-date the GP deposit whereas the breakout lobe on A, and K, L, M, N, and OP

FIGURE 10 | (a) Map of domes K, L and M, and connecting fissures, based on AUV bathymetric map of the area. The white dashed line across K shows the position of the topographic profile given in (b) and the small inset map shows the location of these domes on the caldera rim. The fissures extend east-west across the gap between K and M, to the north of L. They are a few m across and 150–180 m long. The fissures lie within a bench that extends from K to M and has been uplifted by ∼10 m. (b) Northwest (LHS)-southeast (RHS) topographic profile of K, showing the central ridge > 30 m tall and 50 m across with a steep axial cleft more than 20 m deep. (c) Full view of a spine on H. The spine is ∼5 m across.

coherent core at the top of the scarp. (d) Laterally jointed coherent rhyolite in the middle of the scarp. (e) Complex radial joints in coherent rhyolite at the base of the scarp. (f) Schematic cross-section of G based on the pre- and post-eruption bathymetry in 2002 and 2015, and the facies exposed in the scarp.

post-date the GP deposit (**Figure 13**). The relationship between the GP deposit and each of B, D, E, F, G, H, and I is not known.

Some contact relationships between adjacent lavas or domes imply their order of emplacement. On the steep southwestern caldera wall, well-sorted decimeter-sized talus breccia of C overlies less-sorted coarser breccia of B to the northwest and E to the southeast (**Figures 14a–c**). Although these observations imply only the final relationship when the lavas stopped moving, the simplest interpretation is that B, E, and probably D (a small dome between C and E) either preceded or are contemporaneous with C. This finding therefore suggests that all of the units erupted from vents on the southwestern caldera wall (A–E) except the small breakout on A were emplaced before the GP deposit.

FIGURE 13 | Stratigraphic relationships among the Havre 2012 lavas and domes, the giant pumice deposit (GP), ash-lapilli deposit (AL), and ash-lapilli-block deposit (ALB), along the southwestern and southern caldera margin. Note that J is a pre-2002 feature. The pink shape below K, L, and M represents a possible shallow intrusion responsible for the uplifted bench that connects K, L, and M. The small cone beside L post-dates the GP deposit but was not systematically examined. Part of N was extrusive; this part is not covered by the GP deposit; the GP deposit covers the surface above the intrusive part of N.

Overlapping relationships also exist among G, H and I on the southwestern caldera rim. Their morphologies suggest that H overlaps G and I (**Figure 14d**) and is therefore younger than both. However, the relationship between these domes and the GP deposit remains unknown as they are outside of the GP dispersal area.

### DISCUSSION

#### The Chronology of the Havre 2012 Eruption

Using the GP deposit as a marker on the sea floor and the recorded formation of the pumice raft at the sea surface, three main stages can be defined for the Havre 2012 eruption. This approach requires assuming that the GP deposit on the sea floor is the counterpart of the pumice raft recorded in satellite imagery on July 18, 2012, and is valid only if this correlation is correct.

The first stage involved the effusion of A, B, C, D, and E from vents at 1,200 to 1,300 m bsl on the southwestern caldera wall (**Figures 15a,b**). A small dome formed at D while lavas with greater volumes descended the steep caldera wall to the caldera floor at 1,500 m bsl. A and C advanced hundreds of meters farther across the flat surface of the caldera floor. At least 0.05 km3 of rhyolite was erupted at this stage, including C (0.034 km3 ) which is the longest (1.35 km) among the Havre 2012 lavas. The beginning and duration of the first stage are unknown. However, a rhyolitic lava as long as C may have taken months to be fully emplaced. For comparison, the obsidian lava from the 2011 eruption of Cordón-Caulle in Chile took ∼3 months to reach ∼3 km from the source (Tuffen et al., 2013), and was active for another 18 months (Farquharson et al., 2015).

The second stage is defined by the widespread deposition of GP on the sea floor (**Figure 15c**) and is based on the observation that most of A, and all of C are overlain by the GP deposit. The breakout lobe on A formed after deposition of the GP deposit, implying that the initial lava- and dome-producing stage immediately preceded the second stage (rather than being a long time earlier). Assuming the GP deposit correlates with the pumice raft, the second stage began on 18 July, 2012, and lasted less than 24 h (Jutzeler et al., 2014). The dispersal trajectory of the GP deposit and lateral grain size variations within it (Carey et al., 2018) strongly suggest that the source was the same vent that subsequently produced the large dome pair OP on the southern caldera rim. The significant seismicity associated with this stage (Global Volcanism Program, 2012), the short duration, and the large combined volume of the GP deposit and the pumice raft (∼1.3 km3 bulk; 0.3 km3 DRE; Jutzeler et al., 2014; Carey et al., 2018) indicate that this stage involved rapid ascent of newly supplied magma.

The third stage is characterized by the eruption of lava domes K, L, M, N, and OP along the southwestern and southern rim of

FIGURE 15 | Chronology of the Havre 2012 lava and dome eruptions. (a) Havre caldera in 2002 (based on Wright et al., 2006). (b) Effusive eruption of A, B, C, D, and E from vents at 1,200 to 1,300 m bsl on the southwestern caldera wall. A small dome formed at D; lavas A, B, C, E descended the steep caldera wall. F, G, H, and I may have also been erupted during this stage, or else after the GP deposit. (c) Eruption of the GP; providing the correlation of the GP deposit and the pumice raft is correct, then the GP-forming event began on July 18th, 2012. The GP deposit has a northwesterly dispersal area that covers all the caldera floor but only parts of the caldera rim. (d) Effusive eruption of domes K, L, M, N, O, and P and breakout of a small lobe on intracaldera lava A, before October 2012 when Havre was surveyed by R/V Tangaroa. F, G, H, and I may have also been erupted during this stage, or else before the GP deposit. After or during emplacement, part of G, H, and I collapsed, forming a scarp along their northern margins.

the caldera (**Figure 15d**). The breakout lobe on A (**Figures 9a,b** and **Supplementary Figure S10**) was also extruded during this stage, and requires that the earlier part of A (pre-GP) was still sufficiently mobile to undergo farther advance. The vent on the southern caldera margin inferred to have been the source of the GP deposit (Carey et al., 2018) in the second stage now became the vent for dome pair OP which is exceptionally large among the Havre 2012 rhyolite units (∼0.11 km3 ). The ash-lapilli-block deposit is found only around OP and probably underlies, and therefore predates them (Carey et al., 2018). After the majority of dome O was built, a small-volume effusion created P, displacing O sideways so that O now forms an annulus around P. The domes K–N emerged from vents on benches along the caldera rim. Although the aggregate erupted volume of K–N is small (∼0.0063 km3 ), the 1.4-km-long alignment of these domes, and fissures connecting them (**Figure 10**), suggest that an additional volume of rhyolite was intruded as a dyke at the same time.

The timing of the emplacement of F, G, H, and I is not clear because they are beyond the dispersal area of the GP deposit. They could have been emplaced about the same time as A–E,

before the GP (**Figure 15b**), or about the same time as K to OP, after the GP (**Figure 15d**). The northern parts of G, H, and I are truncated by an arcuate scarp (**Figure 5c** and **Supplementary Figure S11**), indicating that they underwent partial collapse after emplacement. The collapse event occurred after the emplacement of the GP deposit because GP clasts on the caldera floor are partly buried by mass-wasting deposits.

The widespread ash-lapilli deposit is found on the GP deposit and on all the lavas and domes except P. However, its stratigraphic and temporal relationships are complex (Murch et al., unpublished). Internal subdivisions appear to record deposition that began following the eruption of the GP deposit, and persisted at least until the emplacement of P. It is inferred to have been produced by explosive and effusive events, as well as intracaldera mass-wasting events, none of which was large enough to have an expression above the sea surface.

The major part of the Havre 2012 eruption is considered to have ended by October 15, 2012, because there was no significant change in the bathymetry data between October 15, 2012 when R/V Tangaroa completed mapping the area and 2015 (MESH voyage). This chronology suggests that the Havre 2012 eruption began and ended with relatively slow lava effusion from multiple vents, which was interrupted by a brief period of higher discharge rate on July 18, 2012 when the pumice raft and GP deposit were erupted.

#### Structurally Controlled Vents

The Havre 2012 lavas and domes were erupted from vents distributed along ∼one third (4.2 km) of the 15-km-long perimeter of the caldera. In detail, the 4.2 km of vents is made up of shorter linear segments (**Supplementary Figure S6**): BCDE, 0.3 km long, trend 120◦ ; FGHI, 0.7 km long, trend 145◦ ; KLMN, 1.1 km long, trend 090◦ . The vent of A is located on the northwestern extension of the BCDE vent alignment, but there is a 250 m gap between A and B. The BCDE and FGHI vent alignments have similar trends but are offset and separated by a 400-m-wide gap and an elevation difference of 200 m. The large dome pair OP does not fall on any of the other 2012 vent alignments but it lies at the northwestern end of a line of older vents with a trend of 150◦ (**Supplementary Figure S6**).

The linear arrangement of the 2012 vents, proximity of adjacent vents and the uniform composition of the rhyolite erupted (Carey et al., 2018) suggest that the 2012 lavas and domes were fed by three structurally controlled dykes (e.g., Fink and Pollard, 1983; Mastin and Pollard, 1988), corresponding to the three distinct vent alignments. The three segments evident from the vents at the surface could be connected at depth, as has been demonstrated in studies of dyke-fed subaerial rhyolite eruptions (e.g., Reches and Fink, 1988). At least for the KLMN vent alignment, there are additional surface features consistent with this interpretation. In particular, domes K and M on the KLMN vent alignment have large spines and clefts parallel to the vent alignment (**Figure 10a**). A bench ∼10 m high and ∼150 m wide extends between K and M (**Figure 10a**); the bench is cut by fissures to the north of L; the bench and the fissures are both parallel to the KLMN vent alignment. The uplift and extension associated with the bench and fissures could result from a shallow syn-eruptive intrusion (pink shape beneath K, L, and M, **Figure 13**). Uplift inferred to be related to a shallow syneruptive intrusion has been identified at Cordón-Caulle (Castro et al., 2016) and at Medicine Lake Volcano (Fink and Anderson, 2017). Given the location of the vent alignments along the caldera perimeter, we conclude that the prime structural control on the 2012 rhyolite feeder dyke(s) was a new or pre-existing calderamargin fault or fault set.

It is also possible to infer that shallow-level dyke emplacement proceeded from the west to south around the perimeter of the caldera during the course of the whole eruption. The basis for this interpretation is the fact that the lavas (A, B, C, E) and dome (D) at the western end predate the domes on the southwestern and southern caldera rim. Furthermore, available compositional data suggest the 2012 Havre rhyolites become progressively more silicic from the west to the south though the overall change is very small (<2 wt.% SiO2; Carey et al., 2018). The western dyke may have tapped slightly deeper (less evolved) and the southern dyke slightly shallower (more evolved) levels of a weakly vertically zoned magma source.

#### Comparison of the Havre 2012 Lavas and Domes With Subaerial Rhyolite Lavas and Domes

Here, we consider whether the Havre 2012 lavas and domes, erupted in ∼900 to 1,300 m of water, differ significantly from subaerial counterparts in terms of dimensions, volumes, aspect ratios, surface features, textures, eruption rates, and structural controls on vents.

The range in diameter (tens of meters to 720 m), maximum thickness (∼20 to 270 m) and volume (0.0003 to 0.11 km3 ) of the ten Havre 2012 domes is comparable to that of subaerial dacitic and rhyolitic domes (e.g., Walker, 1973; Calder et al., 2015). The surface features of the submarine Havre domes are also similar to those of small subaerial domes, particularly those named spiny or Pelean domes, characterized by tall vertical spines (e.g., Fink and Griffiths, 1998; Fink and Anderson, 2000). The overall organization of a coherent (though jagged and spiny) core surrounded by talus is shared by both subaerial domes (e.g., Swanson et al., 1987; Calder et al., 2015) and the Havre submarine domes, and in both settings, the outer part of the core is vesicular (**Figures 11a,b**). Texturally, the Havre domes match the "obsidian dome" category of Calder et al. (2015), given the low abundance of phenocrysts, but the proportion of genuine obsidian (dense glass) within them is unknown.

Being covered almost entirely by breccia, the Havre 2012 lavas resemble subaerial, intermediate to silicic, blocky lavas (Kilburn, 2000; Harris and Rowland, 2015). In both settings, the breccia is autoclastic; in the Havre case, both quench fragmentation and dynamic stressing probably operated. Among the best subaerial examples of silicic, blocky lavas are the Roche Rosse rhyolite lava (Italy; ∼2 km long, 0.03 km3 ; Bullock et al., 2018), the Cordón-Caulle obsidian lava (Chile; ∼3.6 km long, 0.8 km3 ; Tuffen et al., 2013) and rhyolite lavas of Newberry volcano (e.g., Big Obsidian Flow, 1.8 km long, ∼0.13 km3 ; Interlake Obsidian

Flow, 1.8 km long, 0.025 km3 ; MacLeod et al., 1995), all of which have dimensions comparable to the Havre 2012 lavas. The arcuate surface ridges on the Havre lavas, especially those on A (**Figure 8**), appear to be identical to similar ridges on subaerial silicic lavas (commonly referred to as surface folds, compressional ridges or ogives). In subaerial settings, surface ridges form where the crust buckles in response to compression behind a stalled lava-flow front (Fink, 1980). The same explanation appears valid in the case of the arcuate surface ridges on the Havre submarine lavas. Although not common or not widely recognized, subaerial rhyolite lavas may include small breakout lobes emplaced after the main lava advance ceased (e.g., breakouts associated with the 2011–2012 Cordón-Caulle obsidian lava, Tuffen et al., 2013). Among the lavas produced during the Havre 2012 eruption, both A and E include small breakout lobes.

The aspect ratios of the Havre 2012 lavas (**Supplementary Table S1**) fall in the range considered typical of subaerial highviscosity lavas by Walker (1973; <50 and commonly <8). Lava C has an aspect ratio of 47.7, near the high end of the range whereas the other lavas (A, B, E, F) have much lower aspect ratios (5.5–16; **Supplementary Table S1**).

The narrow-tongue segments of the Havre 2012 lavas on the steep caldera wall superficially resemble channelized lavas (e.g., Harris and Rowland, 2015). However, the strongly triangular cross-section shape and very shallow depth of the central depression (**Figures 7a,b**) are important differences, especially with regard to the identification of levees which are an integral part of channelized lavas. Although there are very few exposures of coherent rhyolite on the narrow-tongue segments, it is reasonable to infer that the interior of the ridges comprised molten rhyolite while the lavas were active, because (1) in the cases of A and C, the narrow-tongue segments fed the widelobe segments on the caldera floor, and (2) there is a breakout lobe on the narrow-tongue segment of E (**Figures 9c,d**). In this respect, the narrow-tongue segments appear to have delivered molten rhyolite to the flow front beneath an insulating cover of breccia. This mechanism of propagation strongly resembles that of subaerial a'a and blocky lavas (Kilburn, 1993; Harris and Rowland, 2015).

By assuming that the GP deposit and the observed pumice raft were contemporaneous, and noting that there was no significant topographical change between the bathymetry mapped on October 15, 2012 and March 2015, Carey et al. (2018) concluded that domes K to P were likely erupted within (or less than) 90 days. The minimum mean discharge rate of dome OP (0.11 km3 ) was 14.4 m3 /s whereas the rates for each of K, M, and N were between 0.14 and 0.37 m3 /s.

Recent subaerial eruptions of Chaiten in 2008 and Puyehue Cordón-Caulle in 2011 demonstrated that rhyolitic effusive eruptions can have maximum discharge rates as high as 70 m3 /s (Pallister et al., 2013; Bertin et al., 2015) which surpasses the rate for typical subaerial dacitic dome eruptions such as Mt. St. Helens (∼0.59 m3 /s; Fink et al., 1990), Unzen (∼3.5 m3 /s; Nakada and Fujii, 1993) or Santiaguito (∼1.4 m3 /s; Harris et al., 2003) by more than an order of magnitude. However, the mean discharge rates of Chaiten (∼16 m3 /s) and Puyehue Cordón-Caulle (16.7 m3 /s) were much lower than the maxima, and comparable to the minimum mean discharge rate of Havre dome OP. The rates for other small Havre 2012 domes K, M and N are comparable to those of small subaerial silicic domes.

We have interpreted the linear alignment of the vents for the Havre 2012 lavas and domes to result from structural control, most likely a fault or fault set. Alignment of multiple vents along faults is commonly shown by subaerial rhyolite lava and dome fields. Well-studied examples include Cordón-Caulle (Lara et al., 2004), the Inyo volcanic chain (Sampson, 1987; Miller, 1985), Newberry volcano (MacLeod et al., 1995), South Sister volcano (Scott, 1987) and Kaharoa domes (Leonard et al., 2002) (**Supplementary Figure S12**).

### Controls on Morphology

Comparison of rhyolitic lavas and domes in subaerial and deep submarine (Havre 2012 rhyolitic lavas and domes) settings has shown the two are closely similar. Deep water above vents might be expected to influence morphology, because elevated confining pressure favors retention of volatiles that lowers the magma viscosity (Murase and McBirney, 1973). On the other hand, enhanced cooling, due to the greater heat capacity of water versus air, might be expected to raise the viscosity (Murase and McBirney, 1973). It could be that in the case of Havre 2012, the two influences were balanced or that neither was significant. The latter option appears most plausible: the Havre melts were probably relatively low in dissolved water, given the anhydrous phenocryst assemblage of the lavas and domes, so elevated confining pressure did little to modify the viscosity of the melt. Enhanced cooling probably influenced only the rate of crust development, and once formed, the crust very effectively insulated the interior from any further cooling by seawater.

Because some Havre 2012 lavas and domes were erupted on a steep slope and others onto flat sea floor, we can assess the influence of substrate slope on morphology. Clearly, the influence was profound – narrow tongues formed in all cases where the rhyolite propagated across the steep caldera wall, and either wide lobes or domes formed where the rhyolite propagated across flat sea floor, both within the caldera or on the caldera rim. On the flat sea floor, the contrast between wide lobes and domes was probably primarily controlled by the available magma volume, small volumes producing domes and larger volumes producing wide lobes.

Observations of active subaerial domes and analog experiments of dome growth have shown that the ratio of discharge rate to cooling rate has a strong influence on dome morphology (e.g., Fink and Griffiths, 1998). In this approach, dome morphology can be used to infer discharge rate. The Havre 2012 domes except G are spiny or Pelean domes; this dome morphology is associated with the lowest discharge and highest cooling rates. The discharge rates for domes K, M, and N (between 0.4 and 1.2 m3 /s; Carey et al., 2018) are at the low end of the known range, consistent with the prediction based on morphology alone. However, the effusion rate of the dome pair OP was substantially higher (14.4 m3 /s; Carey et al., 2018). This much higher effusion rate relates mainly to the effusion of O because P is a lot smaller than O. The morphological

interpretation of O is complicated because it was modified when P was emplaced, but it best fits the spiny category of Fink and Griffiths (1998), despite the much higher effusion rate. It could be that application of this morphological approach to the Havre domes requires that the original scaling and dimensional analysis be revisited for the different ambient conditions in the submarine environment.

Among the domes, G is distinctive in being much lower in relief than the other domes and having well-developed arcuate surface ridges rather than spines typical of the other domes (**Figures 6c**, **14d**). This morphology matches the axisymmetric dome category of Fink and Griffiths (1998), associated with relatively high discharge and low cooling rates. This difference in effusion style was presumably viscosity-controlled, the viscosity of the magma that produced G being lower than that of the other domes. Lower viscosity favors higher discharge rate which in turn influences crust development and outflow distance (e.g., Walker, 1973; Fink and Griffiths, 1990). The cause of G rhyolite having a lower viscosity than the rhyolite that formed the other domes has not been identified.

### CONCLUSION

The Havre 2012 eruption produced 15 submarine rhyolite lavas and domes with a collective volume of 0.21 km3 from 14 separate vents. Five vents (A–E) are aligned along the steep caldera wall between 1,200 and 1,300 m bsl. Four of the five vents produced lavas that flowed to the base of caldera wall, two of which (A, C) extended 450–650 m farther across the flat caldera floor. On the caldera rim (960–1,060 m bsl), eight vents produced domes (G–I, K–P) and one vent produced a lava (F) that flowed 1,070 m down the caldera wall. One caldera-rim vent built two domes (OP), the collective volume of which (∼0.11 km3 ) is the largest of all the Havre 2012 lavas and domes.

On the steep caldera wall, the lavas consist of narrow tongues. The narrow-tongue segments of lavas A and C are connected to wide lobes on the caldera floor. Although covered by breccia, the narrow tongues probably fed molten rhyolite to the lava flow fronts and also to small breakout lobes. Well-developed arcuate surface ridges on the lavas are the crests of folds developed in response to compression during flowage. With the exception of G, the surfaces of the domes consist of jagged, spiny coherent rhyolite surrounded by aprons of talus and have morphologies typical of low discharge rates. The low profile of G and its lobate morphology imply higher discharge rates which in turn probably reflect lower magma viscosity. The presence of narrow-tongue morphology on the steeply sloping caldera wall versus domes and wide lobes on flat sea floor demonstrates that substrate slope was a major control on morphology. Additional analysis of the Havre lava and dome morphologies could yield information on discharge rates for the lavas erupted from the caldera-wall vents and the domes that lie outside the dispersal of the GP deposit.

The dome and lavas erupted from vents on the caldera wall are older than the GP deposit but probably only slightly older because a breakout lobe from one of them (A) is younger than the GP deposit. Domes K, L, M, N, and OP on the caldera rim are younger than the GP deposit. For these domes, minimum mean eruption rates from 0.14 to 14.4 m3 /s can be estimated providing correlation of the GP deposit with the pumice raft witnessed on July 18, 2012, is correct. The ages of lava F and domes G, H, and I relative to the GP deposit could not be established because they are located outside the dispersal area of the GP deposit.

The Havre 2012 vents define three alignments close to the caldera margin. The similar magma composition, strong alignment of vents, small separation between adjacent vents and local presence of ground fissures suggest the vents were fed by fault-controlled dykes.

The Havre 2012 lavas and domes are very similar to subaerial rhyolite lavas and domes in terms of dimensions, volumes, aspect ratios, textures, morphology, discharge rate, and structural controls on vent positions.

### AUTHOR CONTRIBUTIONS

FI, JM, RC, AS, and MJ were on the 2015 MESH voyage. FI, JM, and RC led the writing and revision of the article and created the figures. RM conducted a scientific study on one component of this study which has been used in this manuscript. AS contributed to expertise on bathymetry and vehicle data. AS and MJ contributed to discussions, and revision of text and figures.

## FUNDING

This research was funded by an Australian Research Council Postdoctoral fellowship to RJC (DP110102196 and DE150101190), and National Science Foundation grants OCE1357443 and OCE1357216. FI was supported by a Tasmanian Government Postgraduate Award.

### ACKNOWLEDGMENTS

This science would not be possible without the captain, operations teams, marine technicians and crew of the RV Roger Revelle, who we thank for their expert help at sea. We thank the Operations team of the ROV Jason and the AUV Sentry. We particularly thank Dana Yoerger from WHOI for the significant level of assistance on the ship preparing AUV surveys and data processing during and after the MESH voyage. We also thank Sandrin Feig, Karsten Goemann and staff of the Central Science Laboratory at the University of Tasmania. Comments from reviewers Jon Fink and Ilya Bindeman are gratefully acknowledged.

#### SUPPLEMENTARY MATERIAL

The Supplementary Material for this article can be found online at: https://www.frontiersin.org/articles/10.3389/feart.2018. 00147/full#supplementary-material

### REFERENCES

feart-06-00147 October 15, 2018 Time: 19:28 # 19

**Conflict of Interest Statement:** The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Copyright © 2018 Ikegami, McPhie, Carey, Mundana, Soule and Jutzeler. This is an open-access article distributed under the terms of the Creative Commons Attribution License (CC BY). The use, distribution or reproduction in other forums is permitted, provided the original author(s) and the copyright owner(s) are credited and that the original publication in this journal is cited, in accordance with accepted academic practice. No use, distribution or reproduction is permitted which does not comply with these terms.

# A Recent Volcanic Eruption Discovered on the Central Mariana Back-Arc Spreading Center

William W. Chadwick Jr.1 \*, Susan G. Merle2 , Edward T. Baker3 , Sharon L. Walker4 , Joseph A. Resing3 , David A. Butterfield3 , Melissa O. Anderson5,6† , Tamara Baumberger2 and Andra M. Bobbitt2

1 NOAA Pacific Marine Environmental Laboratory, Newport, OR, United States, 2 CIMRS, Oregon State University, Newport, OR, United States, 3 JISAO, University of Washington, Seattle, WA, United States, 4 NOAA Pacific Marine Environmental Laboratory, Seattle, WA, United States, 5 Department of Earth Sciences, University of Ottawa, Ottawa, ON, Canada, 6 GEOMAR, Helmholtz Centre for Ocean Research, Kiel, Germany

#### Edited by:

Nancy Riggs, Northern Arizona University, United States

#### Reviewed by:

James D. L. White, University of Otago, New Zealand Jan Marie Lindsay, University of Auckland, New Zealand

> \*Correspondence: William W. Chadwick Jr.

<EMAIL_ADDRESS>
#### †Present address:

Melissa O. Anderson, Department of Earth Sciences, University of Toronto, Toronto, ON, Canada

#### Specialty section:

This article was submitted to Volcanology, a section of the journal Frontiers in Earth Science

Received: 29 June 2018 Accepted: 02 October 2018 Published: 23 October 2018

#### Citation:

Chadwick WW Jr, Merle SG, Baker ET, Walker SL, Resing JA, Butterfield DA, Anderson MO, Baumberger T and Bobbitt AM (2018) A Recent Volcanic Eruption Discovered on the Central Mariana Back-Arc Spreading Center. Front. Earth Sci. 6:172. doi: 10.3389/feart.2018.00172 Submarine volcanic eruptions are difficult to detect because they are hidden from view at the bottom of the ocean and far from land-based sensors. However, most of Earth's volcanic activity is in the oceans along tectonic plate boundaries, and modern tools of oceanography now allow us to find and study recent eruptions in the deep sea. The first known historical eruption on the Mariana back-arc spreading center was discovered in December 2015 during exploration of the southern back-arc for new hydrothermal vent sites. A water-column survey along the axis of the back-arc showed hydrothermal plumes over the site characterized by low particle concentrations and relatively high reduced chemical anomalies. A dive with the autonomous underwater vehicle Sentry collected high-resolution (1 m) multibeam sonar bathymetry over the site, followed by a near-bottom photographic survey of a smaller area. The photo survey revealed the presence of a pristine, dark, glassy lava flow on the seafloor with no sediment cover. Venting of milky hydrothermal fluid indicated that the lava flow was still warm and therefore very young. A comparison of multibeam sonar bathymetry collected by R/V Falkor in December 2015, to the most recent previous survey of the area by R/V Melville in February 2013, revealed large depth changes in the same area, effectively bracketing the timing of the eruption within a window of less than 3 years. The bathymetric comparison shows the eruption produced a string of lava flows with maximum thicknesses of 40–138 m along a distance of 7.3 km (from latitude 15◦22.30 to 15◦26.30N) between depths of 4050–4450 m bsl (meters below sea level), making this the deepest known historical submarine volcanic eruption on Earth. The cross-axis width of the lava flows is 200–800 m. The Sentry bathymetry shows that the new lava flows are constructed of steep-sided hummocky pillow mounds and are surrounded by older flows with similar morphology. In April and December 2016, two dives were made on the new lava flows by remotely operated vehicles Deep Discoverer and SuBastian. The pillow lavas have many small glassy buds on the steep flanks of the mounds, locally thick accumulations of hydrothermal sediment near the tops of mounds, and small cones of radiating pillows at their summits. The 2015–2016 observations show a rapidly declining hydrothermal system on the lava flows, suggesting that the eruption had occurred only

**26**

months before its discovery in December 2015. The morphology of the pillow lavas is similar to other historical eruption sites, so the greater depth and ambient pressure of this site had no apparent effect on the processes of lava extrusion and emplacement. This study reveals that some segments of the Mariana back-arc have active magmatic systems despite the relatively low spreading rate, and that other eruptions are possible in the near future.

 For example, Rubin et al. (2012) highlighted the fact that only 17 deep [>500 m bsl (meters below sea level)] submarine eruptions were known to have occurred in the last 500 years, compared to 497 known eruptions on land. Historical deep-sea eruptions have been found in a variety of ways, including by distinctive hydrothermal plumes detected during water-column surveys, repeated bathymetric mapping showing depth changes, time-series visual observations by camera or submersible, radiometric dating of young lava flows, remote detection of seismic swarms by hydrophone or seismometers, and rarely by pumice rafts appearing on the ocean surface (Cowen et al., 2004; Dziak et al., 2011; Baker et al., 2012; Rubin et al., 2012; Carey et al., 2018). Only two seamounts have been directly observed erupting, NW Rota-1 in the Mariana arc and West Mata in the NE Lau Basin (probably both long-lived eruptions) (Chadwick et al., 2008; Resing et al., 2011; Embley et al., 2014; Schnur et al., 2017). More recently, a cabled observatory established at Axial Seamount captured an eruption in 2015 (Kelley et al., 2014; Chadwick et al., 2016; Nooner and Chadwick, 2016; Wilcock et al., 2016). Most of these known eruptions have been found on mid-ocean ridges or in submarine volcanic arcs. We know even less about the character and frequency of deep-sea eruptions in back-arc settings.

Recent systematic exploration of the southern Mariana back-arc between 13 and 18.5◦N during research cruises on R/V Falkor in 2015 and 2016 (FK151121 and FK161129) led to the discovery of new hydrothermal vent sites along the back-arc spreading axis (Baker et al., 2017; Butterfield et al., 2018). That work included the collection of new EM302 multibeam sonar bathymetry along the back-arc that permitted a refined mapping of the axis of spreading, and allowed geologic interpretation of the back-arc spreading segments and their tectonic and magmatic character (Anderson et al., 2017). The R/V Falkor cruises also included high-resolution bathymetric mapping of selected sites with the autonomous underwater vehicle (AUV) Sentry in 2015, followed by dives with the remotely operated vehicles (ROV) SuBastian in 2016. In between those, the NOAA Ship Okeanos Explorer made three exploratory dives with the ROV Deep Discoverer (D2).

During the exploration for new vent sites we discovered evidence of a very recent eruption on the axis of the central Mariana back-arc. The only other known historical submarine eruptions on back-arc spreading centers are on the NE Lau Spreading Center (NELSC), including one discovered in November 2008 when hydrothermal event plumes were encountered by chance in the overlying water column during a regional survey (Baker et al., 2011; Rubin et al., 2012), and eruptions near Tafu cone on the NELSC that were recently discovered by before-and-after bathymetric surveys and confirmed by ROV dives (Rubin et al., 2018a,b).

Here, we describe the physical setting of the recent eruption site, constrain the timing of the eruption to within a 3-year period based on before-and-after surveys, show high-resolution bathymetry and ROV visual observations of the young lava flows that provide information about their emplacement, and present observations and sensor data that show a rapidly diminishing hydrothermal system associated with the new lava flows. These results confirm the value of collecting repeated ship-based bathymetry in potentially active submarine volcanic settings for detecting eruptions, illustrate how AUV-based high-resolution bathymetry helps interpret submarine lava flow morphology and emplacement, and reveal that some segments of the Mariana back-arc have active magmatic systems, despite the relatively low spreading rate.

#### Geologic Setting

The Mariana back-arc spreading center is an example of upper plate spreading in an oceanic subduction setting (Fryer, 1995; Stern et al., 2003). Here the Pacific plate subducts beneath the Philippine Sea plate forming the Mariana trench and the active volcanic arc, which is made up of nine islands and over 60 seamounts, a third of which are hydrothermally active (Embley et al., 2007; Baker et al., 2008; Resing et al., 2009). The Mariana back-arc spreading center is located to the west of the volcanic arc and transitions from a relatively shallow ridge in the south (∼3000 m bsl) to a series of deep basins (up to 5000 m bsl) north of 13◦N that are segmented and arranged in an en-echelon pattern (**Figure 1**).

In this paper, we focus on the Mariana back-arc segment centered at 15.5◦N [following the naming scheme of Anderson et al. (2017)], which has a spreading rate of ∼25–40 mm/yr (Kato et al., 2003). This segment is ∼34 km long, with an axial valley that is ∼22–25 km wide, with a maximum depth that ranges from 4650 m bsl in the south to 4350 m bsl in the north (**Figure 2**). The neovolcanic zone is dominated by hummocky volcanic morphology with little or no faulting and has an axial

volcanic ridge that rises up to ∼1050 m above the surrounding seafloor with a shallowest depth of 3820 m bsl near 15.5◦N (Anderson et al., 2017). Anderson et al. (2017) classified this segment as a tectonic segment currently undergoing magmatic extension, where extension is accommodated by the intrusion of dikes, and the morphology is characterized by an axial valley with moderate-to-high relief at the segment center (600–1300 m), a prominent hummocky axial volcanic ridge (800–1050 m) without a central graben, and limited faulting within the axial valley. Anderson et al. (2017) also calculated the volume of neovolcanic material on this segment as a first-order estimate of eruption rates and magma supply, in comparison to the other segments in the Mariana back-arc. Interestingly, the 15.5◦N segment had the second-lowest calculated eruption rate per kilometer (3912 m3 /yr/km) of all the segments in their study area (from 12.7◦N to 18.3◦N). So, in some ways, this segment was the least likely to have hosted a recent eruption. However, the calculated long-term eruption rate is based on its morphology produced by activity over perhaps the last 100,000 years. In contrast, the recent eruption is a manifestation of the present-day (or very

recent) magma supply. Therefore, the recent eruption on this segment is consistent with the characterization that the segment is undergoing recent magmatic rejuvenation, as is the discovery of high-temperature black-smoker vents at the Perseverance vent field located on the along-axis high just 5 km north of the eruption site (**Figure 2**; Baker et al., 2017; Butterfield et al., 2018; Chadwick et al., 2018).

#### Discovery of the Recent Eruption Site at 15.4◦N in December 2015

The search for new hydrothermal vent sites along the southern Mariana back-arc during the FK151121 expedition included first using a towed Conductivity, Temperature, Depth (CTD) instrument package. The CTD was raised and lowered from 10 s to several hundred meters above the seafloor as the ship slowly drove forward, resulting in the CTD producing a saw-tooth path in a vertical 2D profile above the seafloor along the tow track (Baker et al., 2017). Where hydrothermal plumes were

discovered with the CTD, we deployed the Sentry to collect high-resolution bathymetry in those areas, which was used later when ROV dives were made to localize the source of the plume on the seafloor. Hydrothermal plumes can be detected by CTD sensors that measure temperature, light scattering [turbidity (1NTU, Nephelometric Turbidity Units) from hydrothermal particles], and oxidation reduction potential [or ORP (1E), for detecting reduced chemicals such as Fe2+, HS−, H2]. On the Sentry mapping dives these same sensors were housed in a self-contained instrument called a Miniature Autonomous Plume Recorder, or MAPR (Walker et al., 2007). The MAPR data from the Sentry dives were complementary to the CTD data (Walker et al., 2016a; Baker et al., 2017).

CTD tow T15B-06 was conducted along the axis of the back-arc segment centered at 15.5◦N (**Figure 2**), and revealed two plumes with distinctly different character (**Figure 3**). To the south of the highest part of the segment, between 15◦ 23.40 to 15◦ 27.60N a plume was detected with relatively low 1NTU values but high 1E values. Further north, between 15◦ 27.60 to 15◦ 33.00N, the opposite was found in a plume that had relatively high 1NTU values and lower 1E values (**Figure 3**). This was interpreted as evidence for low-temperature diffuse venting in the south and hightemperature focused venting in the north (Baker et al., 2017).

To investigate further, we deployed Sentry for dive 367, which collected multibeam sonar and plume data in a 2- × -3.5-km area at an altitude of 70 m above the bottom, and then conducted a photo survey at a lower altitude of 5 m over a smaller area (500 m × 1000 m) centered on the highest 1E anomaly from the CTD tow (**Figure 4**). We had anticipated that we might photograph some vent animals at a diffuse venting site, but unexpectedly the Sentry photo survey revealed pristine young pillow lavas covering some of the area (**Figures 5a,b**). The young lava flow was dark, glassy, and completely lacking in any sessile animals or even pelagic sediment. A very light dusting of hydrothermal sediment could be seen in a few areas at the tops of pillow mounds (**Figure 5c**), and within one of those areas a cloud of milky hydrothermal fluid was photographed coming out of the flow (**Figure 5d**). The surrounding lava flows are clearly much older with a few cm of sediment completely mantling the lava lobes (**Figure 5e**). On the eastern edge of the photo survey, in low-lying areas below 4300 m bsl that are noticeably more faulted, the older seafloor is locally completely buried by sediment (**Figure 5f**).

The MAPR data collected by Sentry during dive 367 (Walker et al., 2016b) showed multiple 1E anomalies indicative of hydrothermal venting in the western half of the survey area during the mapping at 70 m above bottom (**Figure 4A**). During the photo survey at 5 m altitude, a large 1E anomaly was recorded at the same time and location that the milky vent fluid was photographed (**Figure 4B**). The photograph of active venting and the number of anomalies detected by the MAPR on Sentry suggest that the lava flows were young enough to be still warm and actively cooling (implying the lavas were only months to years old). On the other hand, chemical analysis of water samples from the CTD tow over the young lava flows showed low levels of hydrogen, indicating that the eruption had ended at least days to weeks beforehand. Active submarine eruptions produce high levels of hydrogen from magma-water interaction, which has a short residence time in the water column (McLaughlin-West et al., 1999; Lilley et al., 2003; Baker et al., 2011; Resing et al., 2011; Baumberger et al., 2014). The 1-m resolution bathymetry collected by Sentry during dive 367 shows

FIGURE 3 | Depth cross-section along the axis of the 15.5◦N segment of the Mariana back-arc showing data from CTD tow T15B-06 (from Baker et al., 2017). Gray is seafloor profile. Colors above seafloor show turbid particle plumes as 1NTU anomalies. Colored lines along CTD tow path (gray zigzag lines) show 1E anomalies in millivolts from ORP sensor (see inset for scale). Red dots are confirmed (solid) or inferred (open) hydrothermal sites from Baker et al. (2017). Yellow highlight on the bathymetry profile indicates extent of new lava flows. Note relatively high 1E but low 1NTU anomalies over new lava flows, and high 1NTU and lower 1E anomalies over northern-most hydrothermal site, at the Perseverance vent field with black-smoker chimneys.

that the areas within the young lava flow photographed by Sentry have smooth (unfaulted) hummocky morphology, typical of pillow lava mounds (**Figure 4B**).

To constrain the age of the young lava flow eruption, we compared (post-eruption) ship-based multibeam bathymetry collected by R/V Falkor with its EM302 sonar system on 01 December 2015 [expedition FK151121 (Resing, 2016b)], with the last survey of the area collected by R/V Melville with its EM122 sonar on 14 February 2013 (expedition MV1302a). Comparing the two surveys revealed a string of five areas with maximum depth change from 40 m up to 138 m (**Table 1**), extending over a distance of 7.3 km along the back-arc spreading axis from latitude 15◦ 22.30 to 15◦ 26.30N (**Figure 6**). The cross-axis width of these depth changes was 200–800 m. For convenience, we number the areas of depth change 1 to 5, from north to south (**Table 1**). The third area of depth change includes the area photographed and mapped by Sentry (**Figures 4**, **6**). The timing of the eruption can be bracketed within the 2.8-year period between the two bathymetric surveys. However, the fact that the young flows were still emitting hydrothermal fluids when first photographed suggests that the eruption was probably late within that time period (Baker et al., 2018). A search for anomalous seismicity in the area that might have narrowed the eruption time window further did not uncover any unusual activity1 , so the eruption was not detected remotely.

1Matt Haney, USGS, personal communication, 2015.

of hummocky pillow lava mounds, which commonly accumulate to thicknesses of tens to over 100 m in divergent plate boundary settings (Caress et al., 2012; Yeo et al., 2012; Chadwick et al., 2013; Yeo et al., 2013; Chadwick et al., 2016; Clague et al., 2017). The outlines of the depth changes (and young lava flows) shown in **Figure 6** are minimum areas, since the detection threshold used was a depth difference of 5 m, and therefore any areas where the flow thickness is less than that (including flow margins) are not resolved. For example, observations from later ROV dives show that some of the separate areas of depth change are actually connected by young lavas (described below). Nevertheless, the before-and-after bathymetry can be used to constrain the thickness and area of the lava flows and the volume of lava erupted (**Table 1**). Our analysis shows that the total area of the young flows is 1.81 × 106 m2 and the total eruptive volume was 66.3 × 106 m3 . For comparison, this volume is larger than all the previously documented historical eruptions on the Juan de Fuca and Gorda spreading ridges in the NE Pacific, except for the 2011 and 2015 eruptions at Axial Seamount, which is a hot spot volcano superposed on the Juan de Fuca spreading ridge (Chadwick et al., 1998; Rubin et al., 2012; Yeo et al., 2013; Clague et al., 2017).

The 2013–2015 eruption is the deepest known historical submarine volcanic eruption, and extends over a depth range of 4050 to 4450 m bsl. The post-eruption bathymetry, the photo survey, and the relatively large magnitude of the depth changes between the surveys indicates that the eruption produced a chain

FIGURE 4 | AUV Sentry bathymetry from dive 367 with 1E anomalies from MAPR instrument shown as pink to purple colored lines (see inset for scale) overlain on Sentry track lines (white; widely spaced tracklines were 70 m above bottom). (A) Map of whole survey (modified from similar figure in Baker et al., 2017). Dashed black box is area of Figure 4B. (B) Detailed map of Sentry photo survey area (white track lines; 5 m above bottom). Black outlines show area of depth change between multibeam surveys; red outlines show edges of the 2013–2015 lava flows based on visual observations and interpreted from features in AUV Sentry bathymetry. AUV Sentry bathymetry was re-navigated and then shifted 60 m east and 20 m north to better match ground-truth from ROV dive observations.

The area surveyed during AUV Sentry dive 367 (Resing, 2016a) covered part of the northern half of the chain of lava flows along the back-arc spreading axis (covering the second and third areas of depth change), but did not extend all the way to the northern end where the thickest of the flows was located (**Figures 4**, **6**). Superposing the outlines of the areas of depth change determined by the ship-based bathymetric comparison on the high-resolution Sentry bathymetry shows that all the 1E anomalies recorded by the MAPR on Sentry were located over relatively thick accumulations of lava (**Figures 4**, **6**). The Sentry bathymetry also shows that the hummocky flows produced by this most recent eruption (on the west side of the Sentry survey area) are very similar in morphology to previous eruptions along this segment of the back-arc (on the east side of **Figure 4A**). In fact, it would be impossible to map the boundaries of 2013–2015 lava flows on the AUV bathymetry without the additional information from multibeam depth changes and the visual observations. In general, the young lava flows have a smooth but hummocky morphology in the AUV bathymetry, which shows that they were emplaced from multiple local eruption centers along a N-S fissure system. The Sentry photo survey was fortuitously located spanning the southeastern edge of the third area of depth change (**Figure 4B**). The navigation of AUV Sentry dive 367 bathymetry was reprocessed using mbsystem navadjust software (Caress and Chayes, 2016) based on matching features in overlapping swaths, and then was shifted 60 m east and 20 m north to best match the ship-based EM302 bathymetry. The location of the edge of the young lavas mapped by the Sentry data and the ROV dives (described below) compared to the areas of depth change between bathymetric surveys gives an idea of the uncertainty in the ship-based multibeam comparison method. The lava flow boundary mapped by visual ground truth (red lines in **Figure 4**) is generally within 50–100 m of the depth change boundary (black outlines in **Figure 4**).

TABLE 1 | Estimates of the thicknesses, areas, and volumes of new lava flows, based on depth changes between 2013 and 2015 bathymetric surveys.

∗ The areas and volumes of depth change are minimum estimates of the areas and volumes of the new lava flows, since they do not include the thin margins of the flows.

maximum depth change within each area. Black boxes show area of Figure 4.

### Visual Characterization of the 2013–2015 Lava Flows and Changes Over Time

After the discovery of the recent eruption site in December 2015, two later expeditions with ROVs visited the site in 2016: NOAA ship Okeanos Explorer cruise EX1605L1 with ROV Deep Discoverer (D2) in April 2016, and R/V Falkor cruise FK161129 with ROV SuBastian in December 2016. Each expedition made one ROV dive on the new lava flows, which allowed for visual observations over different parts of the eruption site, sampling of the lava flows, and a search for evidence of any on-going hydrothermal activity.

ROV Deep Discoverer (D2) made dive EX1605L1-09 on 29-30 April 2016 on the northern-most and thickest of the 2013–2015 lava flows (**Figure 7**), for which we only have ship-based bathymetry (∼40 m resolution), because it is north of the Sentry survey. The ROV dive made a north-to-south zig-zag traverse

over three haystack-shaped mounds of young pillow lavas, each 50–100 m high and 200–400 m wide, along a horizontal traverse of ∼750 m, during ∼5.5 h of time on the bottom. The second of the three mounds has the thickest accumulation of pillow lavas, with a maximum depth difference of 138 m, although the ROV track skirted the very thickest part (**Figure 7**). The pillow lavas on the relatively steep sides of the mounds frequently were decorated with an extraordinary number of extremely glassy "buds" or "fingers" of lava (5–10 cm in diameter) extending outward for a few 10s of cm from the main pillow tubes (0.5–1.0 m in diameter) (**Figures 8a,b**). These are similar to the "knobby pillows" described by Ballard and Moore (1977) on the mid-Atlantic ridge, interpreted to be indicative of faster flow rates on the steeper flanks of pillow constructs. The flanks of the mounds had little or no hydrothermal sediment on the lava. In contrast, the tops of the pillow mounds had much broader and flatter pillows and were commonly dusted with low-temperature hydrothermal sediment composed mainly of iron oxy-hydroxide, which locally accumulated into thicker deposits between the lobes (**Figures 8c,d**). The pillow mounds were topped with small conical constructs that were ∼10 m high and ∼5 m in diameter and formed of radiating pillow tubes (**Figures 8d,e**), apparently representing late-stage eruption centers. These are

similar to the "pillowed cones" described in Ballard and Moore (1977).

The north slope of the second, thickest mound was nearly vertical and consisted of a mix of intact and truncated pillows (**Figures 7**, **8f**), with an apron of talus at the bottom consisting of pillow fragments (**Figure 8g**). The talus was likely primary and formed by auto-brecciation during lava flow emplacement and mound construction, rather than due to later tectonism. Beyond the apron of talus, smaller angular glassy fragments were deposited on top of pillow lavas and extended outward for 50–75 m from the base of the cliff (**Figure 8h**), a more distal component of auto-brecciation, also formed by pillow fragments tumbling down the cliff. The dive started and ended within the new lavas and no contacts with older lavas were seen.

During the dive we saw one area of noticeable diffuse hydrothermal venting where we measured a temperature of 7.0◦C with the ROV's temperature probe, well above the ambient temperature of 1.65◦C (**Figure 9a**). Vent endemic species were commonly sighted throughout the dive, including polychaetes (**Figure 9b**), shrimp (**Figure 9c**), and squat lobsters (**Figure 9d**). This is additional evidence that the lava flows were still actively cooling in April 2016, and had hosted diffuse hydrothermal venting long enough to be colonized by mobile vent animals,

FIGURE 8 | Images from ROV Deep Discoverer dive EX1605L1-09 on 29–30 April 2016, which crossed the northern-most and thickest of the 2013–2015 lava flows. See Figure 7 for photo locations. Numbers in parentheses are time of photo in GMT and horizontal scale. (a) Pillow lavas with extremely glassy finger-like buds (03:41:31, 2 m). (b) ROV set up to sample a glassy pillow bud (03:49:32, 4 m). (c) Broader lobes on top of one of the pillow mounds with hydrothermal sediment (00:03:48, 5 m). (d) Conical eruptive center at the top of a mound with radiating pillow tubes and hydrothermal sediment (02:43:00, 6 m). (e) View of ROV illuminating a conical eruptive center (02:40:20, 20 m). (f) Nearly vertical cliff with intact and truncated pillows (01:12:11, 10 m). (g) Primary pillow talus at base of cliff (01:06:06, 3 m). (h) Fragmental deposit on top of pillows beyond the talus apron (01:03:14, 2 m).

but not by sessile species that take longer to colonize. This is consistent with the lava flows only being months to years old when first discovered. The nearest known hydrothermal site is the Perseverance vent field located ∼5 km to the north on the segment high (**Figures 3**, **4**).

Seven months later on 12 December 2016, a second ROV dive was made during expedition FK161129 from R/V Falkor with ROV SuBastian. Dive S45 started 240 m SSE of the end of the previous ROV dive and explored areas further south where high resolution bathymetry had been collected by AUV Sentry a year earlier (**Figure 10**) and where the MAPR on Sentry detected 1E anomalies, indicating hydrothermal activity (**Figure 4A**). The dive started at a depth of 4045 m bsl in young lava with a light dusting of yellow hydrothermal sediment (**Figure 11a**) at the top of a pillow mound at the SW end of the northernmost area of depth change (**Figure 10**). The south side of that pillow mound was a nearly vertical cliff, over 110 m high, that was mantled by almost completely intact young pillow lavas (**Figure 11b**). The pillows on the cliff face were narrow glassy elongated tubes appearing like elephant trunks or the drips of wax

on the outside of a candle (**Figures 11c,d**). It was remarkable how little talus was at the base of the cliff, in strong contrast to the cliff encountered on the previous dive (**Figures 8f–h**), apparently because these lavas were emplaced at a higher extrusion rate so the pillows remained intact even though the lava was flowing down such a steep slope. Extrusion rate would affect the thickness of lava crusts during pillow emplacement, with more rapidly emplaced pillows more likely to have a thinner plastic crust before solidification, rather than a thicker brittle crust that is more likely to result in auto-brecciation. The location of this cliff on the Sentry bathymetry coincides with a navigation artifact, so on the map it appears even more steep than it actually is (**Figure 10**). The cliff on the map is between the northern-most two areas of multibeam depth change, but this ROV dive showed that young lavas are continuous and actually connect the two areas (but the lava must be relatively thin there). On **Figures 4**, **10**, the red outlines show our interpreted mapping of the extent of the young lava flows, based on ROV observations and the Sentry high-resolution bathymetry.

South of the cliff, the ROV track crossed over several low pillow mounds, each 20–25 m high and 70–100 m wide. At the top of the first one, we found another steep-sided pillow cone representing a local eruptive center, covered with a fine dusting of hydrothermal sediment (**Figure 11e**). The high-resolution bathymetry shows that many of the pillow mounds in this area (both young and old) have these pillow cones at the shallowest points (they appear as small dimples on the map). The dive proceeded southward to the top of the highest of the pillow mounds within the second area of multibeam depth change, which is located near the southern edge of the area of depth change, and is topped by a ridge with an E-W orientation (**Figure 10**). Here the ROV turned and headed west, zigzagging along the crest of the E-W pillow ridge and followed it down to the western contact between young and old lavas (**Figures 11g,h**). There, the young lavas were again characterized by larger pillows with many glassy pillow buds. The remainder of the dive continued south onto an older pillow ridge that separates the second and third areas of depth change (**Figure 10**). The older lavas had moderate sediment accumulation and were colonized by sessile animals such as sponges, anemones, and crinoids (**Figures 11g,h**). The dive covered 1.6 km on the bottom in 3.75 h (twice as far as the ROV D2 dive in less time), but the dive had to be terminated early due to deteriorating weather.

While traversing over the younger lavas during ROV SuBastian dive S45, we saw several areas with thick accumulations of yellow hydrothermal sediment (**Figure 11f**), but no visible fluid flow (shimmering water). However, we had a MAPR instrument on the ROV and the data show minor temperature and 1E anomalies in the areas with thick hydrothermal sediment (**Figure 12**) especially on the E-W pillow ridge near the end of the dive, indicating that hydrothermal fluids must still have been seeping out, but so diffusely that they were not visible. Although the two ROV dives were in different (but adjacent) areas, the observations are consistent with a rapidly waning hydrothermal system as the lava flows were cooling in the aftermath of the eruption. Sentry saw robust 1E anomalies 70 m above the lava flows in December 2015, then ROV D2 found only one area of visible venting and elevated temperature on the flows in April 2016, and finally ROV SuBastian did not encounter any visible diffuse flow,

with ROV dive track overlain. (C) Depth profile along ROV dive S45 track line, showing pillow mounds (Sentry bathymetry). Vertical dashed lines indicate major bends

but still detected barely measurable instrumental anomalies near the seafloor. Likewise, CTD tow T15B-06 and vertical cast V15B-06 in 2015 showed the rise height of 1E anomalies was > 400 m above bottom (**Figure 3**), whereas CTD cast V16A-03 in 2016 near the same location had only a weak 1E anomaly at less than 100 m above bottom. In terms of biological colonization, on ROV SuBastian dive S45 we saw some of the same vent-endemic animals as on the previous ROV D2 dive (polychaetes, shrimp, and squat lobsters), but in much fewer numbers, also consistent with a rapidly waning hydrothermal system. All these observations suggest that the eruption occurred relatively late in the 2013–2015 time window constrained by the multibeam sonar surveys. They also imply that the hydrothermal activity from the lava flows was likely temporary and short-lived, consistent with observations of the hydrothermal response after multiple eruptions at Axial Seamount (Chadwick et al., 2013; Baker et al., 2018).

### DISCUSSION

This study was a fortuitous outgrowth of a large-scale exploration of the spreading axis of the southern Mariana back-arc for new hydrothermal vent sites between 13 and 18.2◦N, an area that had not been systematically surveyed previously. The discovery of the new vent sites (Baker et al., 2017) and characterization of their chemistry and biological communities (Tunnicliffe et al., 2017; Butterfield et al., 2018) will fill a knowledge gap to help interpret the biogeography of the region and the connections between geology, chemistry, and chemosynthetic ecosystems (Chadwick et al., 2018). Understanding such links will have important implications for management plans of the Mariana Trench Marine National Monument.

Judging from the relatively low spreading rate of the central Mariana back-arc segments [∼25–40 mm/yr (Kato et al., 2003)], individual eruptions would be expected to be relatively rare

in the dive track.

FIGURE 11 | Images from ROV SuBastian dive S45 on 12 December 2016. See Figure 10 for photo locations. Numbers in parentheses are time of photo in GMT and horizontal scale. (a) Young pillow lava at the top of a mound at the beginning of the dive (01:31:24, 4 m). (b) Nearly vertical cliff mantled by intact young pillow lavas (01:47:42, 15 m). (c) Glassy narrow pillow tubes on the cliff face (01:48:26, 12 m).  (e) Pillow cone atop a low pillow mound dusted with yellow hydrothermal sediment (02:14:58, 8 m). (f) Locally thick accumulation of hydrothermal sediment on a young pillow mound (02:21:05, 10 m). (g) Western contact of the young lavas (glassy in upper left) with surrounding older lavas (with sponge attached at lower right) (04:19:45, 5 m). (h) Surrounding older lavas with anemone attached (04:22:14, 5 m).

(perhaps only every 100 years or more). On the other hand, Baker et al. (2017) point out that back-arc basins can have both enhanced incidence of hydrothermal venting and an enhanced magma supply where they are close to adjacent magmatic arcs, although this back-arc segment is not particularly close to the Mariana arc. In any case, it was surprising to discover fresh new lava flows on the 15.5◦N segment of the Mariana back-arc. There have been relatively few deep-sea (>500 m depth) eruptions documented worldwide (<40), due to the difficulty in detecting them far away from land-based sensor networks (Dziak et al., 2012; Rubin et al., 2012), despite the fact that ∼75% of Earth's volcanic output is in the oceans (Crisp, 1984). That makes finding and characterizing recent deep-sea eruptions rare, valuable, and informative. The investigation of these historical eruption sites provides information on the frequency of eruptions, their volumes, and their chemical and biological impacts in the deep-sea. They also give us a glimpse of the fundamental process of seafloor spreading and ocean crust formation.

It is known from laboratory analog experiments that pillow lavas form at relatively low extrusion rates whereas sheet flows form at higher extrusion rates (Griffiths and Fink, 1992; Gregg and Fink, 1995). More recent studies have interpreted larger constructional volcanic features in terms of extrusion rate. For example, studies at the hotspot-influenced Axial Seamount on the Juan de Fuca ridge in the NE Pacific have used 1-m resolution AUV bathymetry and ROV observations to document a range of morphologies on lava flows erupted in 1998, 2011, and 2015 (Caress et al., 2012; Chadwick et al., 2013, 2016; Clague et al., 2017). Chadwick et al. (2013) introduced the concept of map-scale lava flow morphology to describe features that are on the scale of hundreds of meters, which are discernable in high-resolution AUV bathymetry. They proposed three types: inflated lobate flows, inflated pillow flows, and pillow mounds, mainly distinguished by the extent of a molten lava core within the interior of each flow type during emplacement. "Pillow mounds" were interpreted to be map-scale constructions without a significant molten core. Clague et al. (2017) modified this terminology and used the terms "channelized flows" and "hummocky flows" instead. The "channelized flows" of Clague et al. (2017) include the "inflated lobate flows" and "inflated pillow flows" of Chadwick et al. (2013), and are interpreted as near-vent and distal facies, respectively, of the same flow type. The "hummocky flows" of Clague et al. (2017) are pillow lavas that form mounds, coalesced mounds, or ridges. The main difference to the "pillow mounds" of Chadwick et al. (2013) is that "hummocky flows" also have small to voluminous molten cores, evident from summit collapses, levee-bounded lava ponds, surface tumuli, and off-fissure hummocky flows fed through tubes.

In this paper, we have used the terms "pillow mounds" and "hummocky flows" interchangeably, and we agree with Clague et al. (2017)'s interpretation that these flow types are commonly emplaced with molten cores in their interiors. However, although the "hummocky flows" erupted at Axial Seamount often show more obvious morphologic evidence of a molten interior, the 2013–2015 flows on the Mariana back-arc do not. While there were clearly multiple eruption centers distributed along their lengths and widths, based on the multiple pillow-cones encountered during the ROV dives and that are evident in the Sentry bathymetry, there are no areas of lava drain-out and collapse or the other features described above. This, and the prevalence of locally steep constructional pillow-lava slopes in some places suggest that the Mariana back-arc hummocky flows were emplaced relatively rapidly, perhaps within a time window of only a few days. This interpretation is based on the observation that the thickest hummocky flows emplaced during the 2015 eruption at Axial Seamount, which had levee-bounded lava ponds at their summits as mapped by Clague et al. (2017), were emplaced over a time period of nearly a month, based on the duration of summit deflation (Nooner and Chadwick, 2016) and the explosion-like sounds recorded by a local network of seismometers and hydrophones (Wilcock et al., 2016; Caplan-Auerbach et al., 2017). The Mariana back-arc flows are more consistent with the observations of "lava domes" or pillow mounds on the axis of the southern East Pacific Rise by White et al. (2000) who interpreted that they formed directly over eruptive fissures over time periods of days. The discontinuous distribution of the hummocky flows along-axis on the Mariana backarc may be evidence of pre-eruption focusing of magma within a mildly overpressurized feeder dike as it rose to the surface, similar to the discontinuous pillow mounds documented by Yeo et al. (2013) on the Juan de Fuca and Gorda ridges.

This eruption in the central Mariana back-arc produced a series of hummocky pillow lava mounds along the spreading axis that were comparable in size and morphology to other recent eruption sites at back-arc and mid-ocean ridge spreading

centers, as well as the older lavas at this site. In fact, despite being the deepest historical eruption documented thus far on Earth (>4000 m bsl), the morphology of the pillow lavas is otherwise indistinguishable from other historical eruption sites, so the greater depth and ambient pressure had no apparent effect on the processes of extrusion, flow, volcanic construction, and solidification. However, it was unusual in that it was one of the first discovered in an active back-arc setting. It was also one of the largest historical seafloor eruptions in terms of volume, 66.3 × 106 m3 based on before-and-after bathymetric surveys, but that may also reflect the small number of documented submarine eruptions to date. Smaller eruptions are probably more frequent, and if so larger less frequent events will only become apparent over time (Perfit and Chadwick, 1998). Perhaps most significantly, this eruption, and the discovery of the Perseverance high-temperature vent field 5 km to the north, indicate that there is probably magma currently stored under the axial high of this segment of the Mariana back-arc, and that other eruptions on this segment may be possible in the not-too-distant future.

## CONCLUSIONS

### DATA AVAILABILITY STATEMENT

The data presented in this study are available at the NOAA Centers for Environmental Information (ship bathymetry), the Rolling Deck to Repository (CTD data), the IEDA Marine Geoscience Data System (Sentry AUV and SuBastian ROV data), and ROV video is available on the Schmidt Ocean Institute YouTube channel.

### AUTHOR CONTRIBUTIONS

WC wrote the manuscript and made most of the figures, was co-Chief scientist on the FK151121 and FK161129 expeditions, and directed the ROV dives described here. SM processed and analyzed ship-based and AUV-based multibeam sonar bathymetric data. EB and SW collected and processed the CTD and MAPR data. JR and DB were Chief Scientists on the FK151121 and FK161129 expeditions respectively. MA mapped the young lava flow contact from the Sentry photo survey. TB analyzed CTD water samples for methane and hydrogen on FK151121. AB assisted with GIS data management, volume calculations, and figure preparation.

## FUNDING

We gratefully acknowledge the funding of the NOAA Ocean Exploration and Research (OER) Program and NOAA's Pacific Marine Environmental Laboratory (PMEL) for support of the science at sea and on shore. This research was supported by the PMEL Earth-Ocean Interactions Program, the Cooperative Institute for Marine Resources Studies (CIMRS), and the Joint Institute for the Study of the Atmosphere and Ocean (JISAO) under NOAA Cooperative Agreements NA11OAR4320091 and NA15OAR4320063. JISAO contribution number 2018-0157 and PMEL contribution number 4806.

#### ACKNOWLEDGEMENTS

We thank Patrick Shore and Doug A. Wiens of Washington University in St. Louis for collecting the 2013 multibeam sonar bathymetry during R/V Melville expedition MV1302a. We also thank the Schmidt Ocean Institute for supporting expeditions FK151121 and FK161129 on R/V Falkor, as well as the ship's crews and the teams that operated AUV Sentry and ROV SuBastian, especially expedition leaders Karl Kaiser (Sentry) and Jason Williams (SuBastian). In addition, we also thank NOAA OER for supporting the Okeanos Explorer operations in the Mariana region in 2016, including the ship's crew, science leads at sea, Deborah Glickson and Diva Amon, and the ROV Deep Discoverer team. The manuscript was improved by the helpful comments from reviewers JL and JW.

#### REFERENCES

feart-06-00172 October 22, 2018 Time: 10:49 # 15

**Conflict of Interest Statement:** The authors declare that the research was conducted in the absence of any commercial or financial relationships that could be construed as a potential conflict of interest.

Copyright © 2018 Chadwick, Merle, Baker, Walker, Resing, Butterfield, Anderson, Baumberger and Bobbitt. This is an open-access article distributed under the terms of the Creative Commons Attribution License (CC BY). The use, distribution or reproduction in other forums is permitted, provided the original author(s) and the copyright owner(s) are credited and that the original publication in this journal is cited, in accordance with accepted academic practice. No use, distribution or reproduction is permitted which does not comply with these terms.

# Why Deep-Water Eruptions Are So Different From Subaerial Eruptions

Raymond A. F. Cas1,2 \* and Jack M. Simmons1

1 School of Earth, Atmosphere and Environment (EAE), Monash University, Clayton, VIC, Australia, 2 Centre for Ore Deposit and Earth Sciences (CODES), School of Physical Sciences, University of Tasmania, Hobart, TAS, Australia

Magmas erupted in deep-water environments (>500 m) are subject to physical constraints very different to those for subaerial eruptions, including hydrostatic pressure, bulk modulus, thermal conductivity, heat capacity and the density of water mass, which are generally orders of magnitude greater than for air. Generally, the exsolved volatile content of the erupting magma will be lower because magmas decompress to hydrostatic pressures orders of magnitude greater than atmospheric pressure. At water depths and pressures greater than those equivalent to the critical points of H2O and CO2, exsolved volatiles are supercritical fluids, not gas, and so have limited ability to expand, let alone explosively. Gas overpressures are lower in deep submarine magmas relative to subaerial counterparts, limiting explosive expansion of gas bubbles to shallower waters. Explosive intensity is further minimized by the higher bulk modulus of water, relative to air. Higher retention of volatiles makes subaqueously erupted magmas less viscous, and more prone to fire fountaining eruption style compared with compositionally equivalent subaerial counterparts. The high heat capacity and thermal conductivity of (ambient) water makes effusively (and/or explosively) erupted magmas more prone to rapid cooling and quench fragmentation, producing non-explosive hyaloclastite breccia. Gaseous subaqueous eruption columns and hot water plumes form above both explosive and non-explosive eruptions, and these can entrain pyroclasts and pumice autoclasts upward. The height of such plumes is limited by the water depth and will show different buoyancy, dynamics, and height and dispersal capacity compared with subaerial eruption columns. Water ingress and condensation erosion of gas bubbles will be major factors in controlling column dynamics. Autoclasts and pyroclasts with an initial bulk density less than water can rise buoyantly, irrespective of plume buoyancy, which they cannot do in the atmosphere. Dispersal and sedimentation of clasts in water is affected by the rate at which buoyant clasts become water-logged and sink, and by wind, waves, and oceanic currents, which can produce very circuitous dispersal patterns in floating pumice rafts. Floating pumice can abrade by frictional interaction with neighbors in a floating raft, and generate in transit, post-eruptive ash fallout unrelated to explosive activity or quench fragmentation.

Keywords: submarine eruptions, hydrostatic pressure, bulk modulus, limited volatile exsolution, supercritical fluid, magma properties, pumice raft dispersal

#### Edited by:

Carles Soriano, Instituto de Ciencias de la Tierra Jaume Almera (ICTJA), Spain

#### Reviewed by:

Rodney Allen, Volcanic Resources AB, Sweden Margherita Polacci, The University of Manchester, United Kingdom

> \*Correspondence: Raymond A. F. Cas<EMAIL_ADDRESS>
#### Specialty section:

This article was submitted to Volcanology, a section of the journal Frontiers in Earth Science

Received: 29 June 2018 Accepted: 22 October 2018 Published: 20 November 2018

#### Citation:

Cas RAF and Simmons JM (2018) Why Deep-Water Eruptions Are So Different From Subaerial Eruptions. Front. Earth Sci. 6:198. doi: 10.3389/feart.2018.00198

## INTRODUCTION

feart-06-00198 November 20, 2018 Time: 12:55 # 2

Understanding the physical processes controlling the dynamics, style and intensity of volcanic eruptions has historically relied upon observations of recent subaerial eruptions (e.g., 1980 Mt St Helens, Lipman and Mullineaux, 1981; 1991 Mt Pinatubo, Newhall and Punongbayan, 1996; 1982 – present Hawai'i, Poland et al., 2014; Soufriere Hills, Montserrat, 1995–2012, Druitt and Kokelaar, 2002) and well-preserved subaerial eruption deposits (e.g., Askja, Carey et al., 2010; Colli Albano, Giordano et al., 2010; Vesuvius, Shea et al., 2011; Santorini, Druitt et al., 1999; Simmons et al., 2016; Tenerife, Marti and Geyer 2009, Edgar et al., 2017). However, volcanism on Earth principally occurs in the submarine realm (at mid-ocean ridge systems, intraplate hotspots/seamounts, oceanic plateaus, oceanic volcanic arcs), under eruptive conditions distinct from atmospheric or subaerial vent settings (McBirney, 1963; Cas, 1992; Head and Wilson, 2003; White et al., 2003, 2015a,b; Wohletz et al., 2013; Cas and Giordano, 2014; Carey et al., 2018; Manga et al., 2018). The heat capacity, thermal conductivity, density, viscosity, pressure gradient and bulk modulus of (sea)water, in particular, are each generally orders of magnitude larger than the properties of the atmosphere at subaerial vents (**Table 1**), which will greatly affect eruption processes.

Although study of both modern and ancient submarine volcanic settings and successions have benefits, they also have limitations as to how much information can be accessed, documented and interpreted (**Table 1**). In ancient deep-water volcanic successions, the major limitation is the lack of understanding of the actual water depth at the time and place of eruption and deposition. As we will see, water depth determines the ambient hydrostatic pressure, and as a result magma properties, eruption style and the deposit characteristics. In modern submarine settings, even if the water depth and hydrostatic pressure are known, unless the physics of the processes at those water depths are carefully evaluated and understood, mistakes in interpreting the eruption processes and origins of deposits are commonly made, including erroneous assignation of some subaerial eruption styles to deep water settings.

In this review, we will outline the physical properties and effects of an ambient water mass on magma properties and eruption dynamics in deep-water environments (>500 m), discuss the deposits or products of subaqueous eruptions from observations in modern and ancient deep-water seafloor settings, and highlight the differences between subaqueous and subaerial eruptions. There are few papers which have attempted this to date (e.g., McBirney, 1963; Cas and Wright, 1987; Head and Wilson, 2003; White et al., 2003; Cas and Giordano, 2014). This current paper represents a significant and comprehensive update.

#### EFFECTS OF HYDROSTATIC PRESSURE ON SUBMARINE ERUPTIONS AND MAGMA PROPERTIES

Magmas erupting from vents in subaerial settings decompress to atmospheric pressures of approximately 0.1 MPa (1 bar, at sea level). This contrasts to magmas erupting on the sea floor, which decompress to the hydrostatic pressure corresponding to the

Ancient submarine successions Modern seafloor settings ADVANTAGES • May preserve stratigraphy (including contact relationships) that provides a 4-D perspective on the evolution of a volcanic center. • Depositional textures may be well preserved. • Deposits are usually accessible. • Fieldwork is relatively inexpensive. • Requires minimal equipment. • Sampling of volcanic material is easy. • Preserve modern submarine edifices. • Preserve deposits without significant reworking or erosion. • Physical setting well constrained (e.g., water depth, topography, tectonic setting). LIMITATIONS • Outcrop is often discontinuous. • Field work can be time consuming if significant mapping is required. • The possible effects of uplift, deformation, hydrothermal alteration, and erosion need to be considered. • Physical setting is NOT well constrained (e.g., water depth, topography etc.). • Limited access to the deep-sea and volcanic stratigraphy. • Very time consuming. • Research is very expensive. • Requires technology to visualize volcanic forms and sample volcanic material.

• Regional context may not be well understood.

TABLE 1 | Advantages and limitations of studies of ancient submarine successions and modern seafloor settings.

water depth of the seafloor at the location of the vent. Hydrostatic pressure, PHYD, is calculated as follows:

$$P\_{\rm HYD} = \text{g}\rho d \tag{1}$$

where g is gravitational acceleration (9.8 m s−2 ), ρ is the density of water (1000 kg m−3 for fresh water and 1020–1030 kg m−3 for sea water), and d is the depth or thickness of the overlying water column (in meters; **Figure 1** and **Table 2**). Hydrostatic pressure thus increases at a rate of ∼9.8 MPa km−1 of water depth, which is ∼10,000× greater than the atmospheric pressure gradient of about −0.009 MPa km−1 with altitude (**Table 2**). At a water depth of 100 m, the hydrostatic pressure is 1 MPa (10× atmospheric pressure, AP), at 500 m, 5 MPa (50× AP), at 1000 m, 10 MPa (100× AP), at 1500 m, 15 MPa (150× AP), at 2000 m, 20 MPa (200× AP; **Figure 1**), and so on. The increasing pressure with increasing water depth has the effect of modifying the physical properties of the erupting magmas, particularly the timing of volatile saturation and exsolution, the state of exsolved volatiles, and magma viscosity (McBirney, 1963; Wallace and Anderson, 2000; Wallace et al., 2015).

#### Effects of Hydrostatic Pressure on Volatile Saturation and Exsolution in Magmas

The saturation pressure of volatile species in magmas, particularly H2O and CO2, is strongly dependent on the initial volatile content, magma composition and properties, and the confining

pressure (**Figure 2**; Wallace et al., 2015). If the volatile content is sufficiently high, volatiles can exsolve from a melt at any pressure/depth in the crust (**Figure 2**; Wallace et al., 2015). When the confining pressure, whether it be hydrostatic, magmastatic or lithostatic, or a combination of these, exceeds the volatile saturation pressure, which is volatile content dependent, volatile exsolution is prevented. For example, a rhyolitic magma containing 4 wt. % H2O and a basaltic magma containing 3 wt. % H2O rising through continental crust in subaerial settings are saturated at confining pressures of ∼100 MPa. Assuming a typical continental lithostatic pressure gradient of 24.5–25 MPa km−1 (**Figure 2**, Wallace et al., 2015), this is equivalent to 4 km depth in continental crust (**Figures 1**, **2**). In submerged crust, the confining pressures will be the sum of lithostatic pressure and the hydrostatic pressure of the water column at the eruption water depth (add ∼9.8 MPa per kilometer of water depth), and the depth in the crust below the seafloor at which volatile exsolution occurs will be:

$$D\_{\rm CR} = (P\_{\rm sat} - P\_{\rm HYD}) / L\_{\rm gr} \tag{2}$$

where DCR is the depth in the crust at which exsolution occurs in meters, Psat is the volatile saturation pressure (MPa), PHYD (MPa) is the hydrostatic pressure at the depth of the vent in the water body, and Lgr is the lithostatic pressure gradient (MPa km−1 ) for submerged continental, arc, or oceanic crust. For example, at a water depth of 1 km, rhyolite magma with 4 wt. % dissolved H2O, or a basalt magma with 3 wt. % dissolved H2O, becomes water saturated at crustal pressures of 90.2 MPa (=100 MPa–9.8 MPa), equivalent to a submarine crustal depth of ∼3.6 km. At 2 km water depth, with a hydrostatic pressure of 19.6 MPa, the lithostatic pressure component is 80.4 MPa, equivalent to a volatile saturation depth in the crust of ∼3.2 km, and at 3 km water depth the volatile saturation pressure equivalent depth in the crust is ∼2.8 km.

Equally importantly, the exsolution of volatiles will cease earlier in submarine settings compared to their subaerial counterpart because the erupting magma decompresses to higher ambient eruption pressures with increasing vent water depths. Consequently, less of the dissolved volatile component will exsolve, meaning a smaller fraction of volatiles is potentially available to drive explosive eruptions in submarine settings. This principle also applies to sub-glacial settings and other planetary bodies with dense atmospheres. For example, on Venus, the atmospheric pressure is 9.2 MPa (Taylor, 2010; Airey et al., 2015), which is equivalent to a depth of ∼1 km in the Earth's oceans.

Geochemical compositional data and pre-eruptive volatile content can be used to determine degrees of exsolution and residual H2O volatile content in magmas rising and decompressing in a conduit at various pressures, using the modeling software CONFORT 15 (Campagnola et al., 2016; **Figure 3**). Using data from Carey et al. (2018) and Manga et al. (2018) for rhyolite samples erupted at water depths of 900 m at the wholly submarine Havre volcano in the Kermadec arc, Southwest Pacific in 2012 (9 MPa confining hydrostatic pressure; pre-eruptive water content of 5.8 wt%; eruption temperature of ∼850◦C), we calculate that the residual water content after

#### TABLE 2 | Physical properties of water and air (from Cas et al., in press).

exsolution at 900 m water depth, and assuming ascent through a 20 m diameter cylindrical conduit, is about one quarter the original magmatic water content (**Figure 3B**).

#### Confining (Hydrostatic) Pressure and the Physical State and Specific Volume of Exsolved Magmatic Volatiles, and Effects on Eruption Styles

Even high levels of vesiculation in erupting subaqueous magmas do not ensure an explosive eruption because the physical state (vapor/gas versus supercritical fluid versus liquid) and the specific volume of exsolved volatiles in magma in conduits or lavas, and of volatile bubbles released into water masses is directly related to the confining pressure (McBirney, 1963; Wallace and Anderson, 2000; Cas and Giordano, 2014; Wallace et al., 2015). The transition between liquid and vapor states occurs at the critical point pressure and temperature of a particular volatile species. At and above the critical point, the fluid is supercritical, which is almost incompressible, and the liquid and vapor states are indistinguishable. The critical point pressure and temperature for fresh H2O (magmatic volatiles) are ∼22 MPa and 374◦C, respectively, and for CO2 they are ∼7.8 MPa and 31◦C, respectively.

For magma in a closed conduit in continental or arc crust under lithostatic pressure, exsolved magmatic H2O bubbles are supercritical at depths >900 m (i.e., the lithostatic pressure is 22 MPa at ∼900 m in the crust). In oceanic crust, the equivalent supercritical pressure depth is ∼750 m, due to the higher lithostatic pressure. In comparison, the supercritical depth of exsolved magmatic H2O in an open column of silicic and basaltic magma in continental crust is ∼980 m and in oceanic crust, ∼830 m. In an open body of water, the water depth that coincides with the supercritical pressure for exsolving magmatic H2O (fresh) is 2200 m, meaning that H2O exsolving in an erupting magma, or released from a vent as fluid bubbles at those depths, is a supercritical fluid. The equivalent critical pressure depth for CO2 is ∼800 m.

Supercritical fluids are dense (∼322 kg m−3 ; Pioro and Mokry, 2011), almost incompressible, and volumetrically limited compared with their gaseous form, and therefore cannot expand explosively (**Figure 4**; McBirney, 1963). If magma

rises to confining pressures less than the critical pressure, the supercritical fluid transforms through boiling to dense steam (∼10 kg m−3 ), leading to a slight increase in specific volume, which then increases exponentially at lower pressures as dense steam transforms into low density steam or gas (∼1 kg m−3 ; **Figure 4**). At P < 10 MPa (equivalent to 1 km water depth), the rate of change in specific volume and abundance of exsolved water bubbles in magma increases dramatically, which reduces the bulk density, but dramatically increases the bulk vesiculated volume, buoyancy and ascent rates of the magma in the crust. The vesiculated magma then has the potential to fragment explosively (**Figure 4**; McBirney, 1963).

Volatile bubbles grow in magma if the exsolved gas bubble (over)pressure, 1P, exceeds the sum of the yield strength of

coinciding with the conditions for the strain rate criterion that would cause explosive fragmentation. In each modeled simulation, a fraction of the available initial volatile content remains dissolved in the magma (top axis) at the point of fragmentation and the depth of fragmentation increases as a function of increasing initial

volatile content. Note, strain rate driven fragmentation only occurs at depths of <500 m, even if the initial magmatic water content was 8 wt. %.

the magma, ys (which for most magmas is several pascals to kilopascals), and the ambient pressure, P (atmospheric, hydrostatic, glaciostatic, lithostatic):

(Modified from McBirney, 1963, Figure 2, by annotating fields and labels; permission granted by Springer Nature License<PHONE_NUMBER>103).

$$
\Delta P \searrow \ y\_s + \text{ ambient } P \tag{3}
$$

Where the yield strength is given by:

$$
\eta\_s = \sigma - \eta \,\eta^n \tag{4}
$$

where σ is the total shear stress applied, η is the viscosity, γ is the strain rate, and n is a constant that is <1 for pseudo-plastic behavior and = 1 for Bingham behavior.

However, vesiculation alone does not ensure that explosive fragmentation, involving brittle bubble wall rupturing, will occur. This can only occur if the gas overpressure, 1P, exceeds the tensile strength, τ, of the magma (<5–6 MPa) and the ambient pressure:

$$
\Delta P > \text{ } \text{ } + \text{ } \\
ambient \text{ } P \text{ } \tag{5}
$$

Papale (1999) recognized that magmas can also fragment during explosive eruptions if the buoyancy driven strain rate that a vesiculating magma is subject to in the conduit exceeds a critical level whereby brittle failure occurs. This criterion for fragmentation in terms of strain rate, γ, and viscosity, η, is based on Maxwell's Law, as follows (Papale, 1999):

$$
\gamma > kG\eta^{-1} \tag{6}
$$