Understanding of Basaltic Eruption Dynamics and Mechanisms: Effusive and Explosive Eruptions In

Understanding of Basaltic Eruption Dynamics and Mechanisms: Effusive and Explosive Eruptions in

Hawaiʻi

by Samuel James Holt School of Natural Sciences - CODES

Submitted in fulfilmentDoctor of the of Philosophyrequirements for the degree of University of Tasmania, April 2018 ii

Declaration of Originality

This thesis contains no material which has been accepted for a degree or diploma by the University or any other institution, except by way of background information and duly acknowledged in the thesis, and to the best of my knowledge and belief no material previously published or written by another per- son except where due acknowledgement is made in the text of the thesis, nor does the thesis contain any material that infringes copyright.

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This thesis may be made available for loan and limited copying and communication in accordance with the Copyright Act 1968.

Statement regarding published work contained in thesis

The publishers of the papers comprising Chapter 2 the copyright for that content and access to the material should be sought from the respective journals. The remaining non-published content of the thesis may be made available for loan and limited copying and communication in accordance with the Copy- right Act 1968.

Samuel J. Holt April 2018 iii

Statement of Co-Authorship

The following people and institutions contributed to the publication of work undertaken as part of this thesis:

Samuel Holt, School of Natural Sciences, CODES – Candidate Rebecca Carey, University of Tasmania – Author 1 Jocelyn McPhie, University of Tasmania – Author 2 Bruce Houghton, University of Hawaiʻi, Manoa – Author 3 Tim Orr, USGS Alaska Volcano Observatory – Author 4 Sandrin Feig, Central Science Laboratory, UTas – Author 5 Jacopo Taddeuci, Istituto Nazionale Geofisica e Vulcanologia –Author 6

Author details and their roles:

Paper 1, Eruption and fountaining dynamics of selected 1985–1986 high fountaining episodes at Kīlauea volcano, Hawaiʻi. Located in chapter 2 Candidate was the primary author and with author 1 and author 2 contributed to the conception and design of the research project as well as reviewing major parts of the paper. All analyses and writing were carried out by the candidate. Author 3 contributed to the conception and design of the project as well as contributing previously collected field samples for analysis. Author 4 provided academic insights and contributed to the review process of the paper as well as providing access to Hawaiian Volcano Observatory databases. Author 5 provided assistance in collecting and interpreting SEM analysis of samples.

Paper 2, Rootless littoral cones in Hawaiʻi formed by sustained, confined mixing of lava and sea-water. Located in chapter 3 Candidate was the primary author and with author 1 and author 2 contributed to the conception and design of the research project as well as reviewing major parts of the paper. All analyses and writing were carried out by the candidate.

Paper 3, Passive outgassing processes of the lava lake at Kīlauea Volcano, Hawaiʻi. Located in chapter 3 Candidate was the primary author and with author 1, 2, 3 and 6 contributed to the conception and design of the research project as well as reviewing major parts of the paper. All analyses and writing were carried out by the candidate. Author 3 and Author 6 contributed previously collected data for analysis. Author 4 provided academic insights and contributed to the review process of the paper as well as providing access to Hawaiian Volcano Observatory databases. iv

We the undersigned agree with the above stated “proportion of work undertaken” for each of the above published (or submitted) peer-reviewed manuscripts contributing to this thesis: Abstract Hawaiʻi is viewed as the “type example” for many fundamental basaltic volcanic phenomena (Garcia, 2017). The extensive geophysical and visual monitoring network maintained by the Hawaiian Volcano Observatory (HVO) (Guffanti et al., 2010) allows for numerous basaltic eruption styles and their products to be described in great detail. Hawaiian volcanoes are famous for effusive eruptions and they are mainly built of pahoehoe and a’a lavas. However, two Hawaiian volcanoes, Kīlauea and Mauna Loa, have been the location of numerous examples, both historical and modern, of other low intensity basaltic eruptive styles. These include Hawaiian-style fountaining, persistent lava lake activity and rootless littoral cone eruptions (Heliker et al., 2003; Jaggar and Finch, 1924; Moore and Ault, 1965; Orr, 2008; Patrick et al., 2016). The products of such activity are rarely preserved in the rock record, and yet they are a common and typical part of these volcanoes’ long-term behaviour and provide valuable constraints on fundamental dynamic processes. Three case studies of such ephemeral eruptive activity, each focussed on a different phenomenon, are presented in this thesis: five episodes of Hawaiian-style fountaining in the 1983-present eruption of Puʻu ʻŌʻō on Kīlauea, the 2010- 2018 stable Halemaʻumaʻu lava lake at the summit of Kīlauea, and the rootless littoral cones at Puʻu Kī and ʻAuʻau Point on Mauna Loa. These studies provide a better understanding of the full spectrum of activity typical of basaltic volcanoes. Tephra from the early Hawaiian fountaining episodes of the ongoing eruption of Puʻu ʻŌʻō in the East Rift Zone (ERZ) of Kīlauea provides an opportunity to study the vesicle microtextures of pyroclasts erupted from a single vent over a prolonged period of time. The results of microtextural analysis of pyroclasts from five of Puʻu ʻŌʻō’s high (>200 m) Hawaiian fountaining episodes (episodes 32, 37, 40, 44 and 45) erupted during 1985–1986 are reported. The aims of the research are to constrain the parameters that led to large variations in fountain height at Puʻu ʻŌʻo, and the extent to which pyroclast residence times in the fountain modified microtextures. The results confirm the finding of Stovall et al. (2011, 2012) that pyroclasts from a single Hawaiian fountain can vary greatly in texture (from bubbly to foamy) and have vesicle

m volume densities (N v) and vesicle to melt ratios (VG/VL) that vary by an order of magnitude. vi

Abstract -3 m 6 Pyroclast microtextures can range from high density (typically >500 kg m ), high N v (2.2×10

6 -3 m to 4.4×10 ), low VG/VL (2.06 to 4.65) bubbly textures to low density (<450 kg m ), low N v

5 5 (2.5×10 to 6.1×10 ), high VG/VL (5.49 to 11.05) foamy textures due to extensive growth and coalescence of vesicles within the fountain after fragmentation. Due to the extensive modification by post-fragmentation vesiculation processes pyroclasts that have primary textures re-

m main elusive and, as a result, it is not possible based on N v to quantitatively describe the fluid dynamics of the magma in the shallow conduit in such a way that allows us to make detailed comparisons between distinct fountaining episodes at a single vent. These data show good correlation between Δ(VG/VL) and peak fountain height, which implies that Δ(VG/VL) can be used as a proxy for peak fountain height in the absence of direct visual observations of the eruption. Two rootless littoral cones, ʻAuʻau Point and Puʻu Kī South Cone, located on the southwestern flank of Mauna Loa, Hawaiʻi have wholly unique cone geometries and cone-forming lithofacies when compared to other examples of historic rootless littoral cone activity elsewhere on Mauna Loa and on Kīlauea. Both cones have, or intitally had, circular geometries and are built from alternating clastogenic lavas and thermally oxidised, slightly welded lapilli and bombs, all of which are deposit features typically associated with ‘dry’ magmatic styles of volcanism. They differ greatly from other examples of rootless littoral activity in Hawaiʻi defined by half-cone deposits consisting predominantly of thinly planar-bedded to cross-bedded, bimodal deposits composed mainly of ash (70-90%) and minor lapilli and bombs. Both ʻAuʻau Point and Puʻu Kī South Cone were formed by prolonged periods of confined mixing (Mattox and Mangan, 1997) of lava and water-saturated substrate driving a cyclic eruption of littoral lava fountaining and bubble burst activity. Despite being formed by confined mixing, these two rootless littoral cones differ from examples in Iceland of rootless cones formed by confined mixing where lavas flowed over water-saturated substrates (Fagents and Thordarson, 2007; Hamilton et al., 2017), highlighting a critical difference in the nature of the molten fuel coolant interactions that drove the rootless eruptions in each case. The identification of rootless cones that have formed from phreatomagmatic activity but have deposits that appear to be of ‘dry’ magmatic in origin (thermally oxidised, variably welded lapilli and bombs), such as Puʻu Kī South Cone and ʻAuʻau Point, casts doubt on the diagnostic features that are used to distinguish between phreatomagmatic and ‘dry’ magmatic activity, and presents a specific challenge for volcanologists working in the ancient rock record. vii

AbstractThe recent summit eruption of Kīlauea began in 2008 and from 2010 to it's demise in May 2018 consisted of a stable lava lake within the Overlook Crater within Halemaʻumaʻu. The physical and chemical processes that occured at the surface of the lava lake were directly related to processes that occured throughout the entire magmatic system of Kīlauea Volcano. In this study, a combined dataset of high-speed thermal infrared videos and near-infra-red time-lapse sequences is used to define and map the spatial distribution of the passive outgassing processes that took place on the surface of the Halemaʻumaʻu lava lake from the 1st to the 6th of December 2013. Observed outgassing features include both boundary and intra-plate bubble bursts, spot vents and surface doming, all of which are different surface manifestations of buoyant gas slugs impacting on the base of the viscoelastic crust of the lava lake surface. At this time, gas slugs arriving at the lake surface were localised in diffuse ring geometry at a fixed point on the lake surface which was not affected by lateral movement of the surface crust. The lateral movement of the surface crust is commonly attributed to convection of magma at deeper levels in the lake. Therefore, this geometry is consistent with the formation of gas slugs at depths below where the lava is mechanically coupled to the surface flow regime. The gas slugs then likely rose through the lava lake decoupled from internal convection. These observations provide basic, qualitative constraints on the internal processes of the lava lake at Halemaʻumaʻu as well as the foundation for further quantitative studies of the internal convection regimes of stable lava lakes at Kīlauea and elsewhere around the globe. Acknowledgments Firstly I would like to thank my supervisors, Rebecca Carey and Jocelyn McPhie. I will always be grateful that you welcomed me down at UTas. Your knowledge, guidance and willingness to teach has been integral in my success during my PhD. Thank you. Further thanks go to my collaborators Bruce Houghton, Don Swanson, Tim Orr, Matt Patrick and Jacopo Taddeucci for their invaluable academic input as well as providing me with the once in a lifetime opportunity to visit Kīlauea Volcano. I would also like to acknowledge Sandrin Feig, Karsten Goemann and Thomas Rodemann at the CSL for their tireless help.

A special thank you must go out to the administrative staff at CODES, specifically (in no par- ticular order) Jane Higgins, Rose Pongratz, Karen Huizing, Helen Scott and Claire Rutherford for their help with the simple but yet seemingly endless administrative problems you encounter as a PhD candidate.

Thank you to all my friends who made Hobart a magical place to live. The life-long friendships I have made is truly what I am most proud of about my time in Hobart.

Thank you Mum and Dad for your unwavering support, endless motivation and exceptional proof reading.

Most importantly, the completion of my PhD would not have been possible if not for the endless love, patience and support given to me by my wife Claire Thomas. I am still not sure how I managed to convince you to move with me to Tasmania in the first place and without you I would never have had the motivation and persistence required to complete my PhD. Table of Contents 1. Introduction 17 1.1 Project Aims 18 1.2 Volcanological Setting 20 1.3 The Geology of Kīlauea Volcano 20 1.3.1 Structure of Kīlauea Volcano 20 1.3.2 The Magmatic Storage System of Kīlauea Volcano 23 1.3.3 1983-Current Activity at Kīlauea Volcano 26 1.4 The Geology of Mauna Loa Volcano 28 1.4.1 Structure of Mauna Loa Volcano 28 1.4.2 Geology and Historic Eruptions of the Southwest Rift Zone of Mauna Loa 28

2. Eruption and fountaining dynamics of selected 1985–1986 high fountaining episodes at Kīlauea volcano, Hawaiʻi. 31 2.1 Introduction 31 2.2 Eruption history of Puʻu ʻŌʻō 33 2.2.1 Studied eruption episodes 34 2.3 Sampling and analytical methods 34 2.3.1 Field work and sample selection 34 2.3.2 Analytical methods 36 2.4 Results 37 2.4.1 Juvenile clast density and vesicularity 37 2.4.2 Microtextural analysis 39 2.4.2.1 Qualitative results 39 2.4.2.2 Quantitative results 39 2.5 Discussion 43 2.5.1 Determining syn-eruptive vesicle textures 43

m 2.5.2 Relationship between vesicle number density (N v) and fountain height 43

2.5.3 Relationship between vesicle-to-melt ratio (VG/VL) and fountain x

Table of Contents height 47 2.5.3.1 Episodes 37, 44 and 45 47 2.5.3.2 Episode 32 48 2.5.3.3 Episode 40 49 2.5.4 Quantitative microtextural signatures as a proxy for fountain height 49 2.6 Conclusions 50

3. Rootless littoral cones in Hawaiʻi formed by sustained, confined mixing of lava and sea-water. 52 3.1 Introduction 52 3.2 Historic rootless littoral cones on Hawaiʻi 53 3.3 Comparison with Icelandic rootless cone groups 56 3.4 Field observations of Rootless littoral cones at Puʻu Kī and ʻAuʻau Point 57 3.4.1 Puʻu Kī South Cone 57 3.4.2 ʻAuʻau Point 59 3.5 Interpretation of deposit characteristics of Puʻu Kī South Cone and ʻAuʻau Point 61 3.6 The formation of Puʻu Kī South Cone and ʻAuʻau Point 62 3.7 Key differences between Puʻu Kī South Cone and ʻAuʻau Point and Icelandic rootless cones 63 3.8 Difficulties in discriminating ‘dry’ deposits from phreatomagmatic deposits in Hawaiʻi based on deposit observations 65 3.9 Conclusions 66

4. Passive outgassing processes of the lava lake at Kīlauea Volcano, Hawaiʻi. 67 4.1 Introduction 67 4.2 Lava lake activity at Halemaʻumaʻu, Kīlauea volcano 68 4.2.1 Summary of the formation and evolution of Kīlauea’s lava lake 68 4.2.2 Summary of gas-pistoning cycles and outgassing regimes of Kīlauea lava lake 70 4.2.3 Summary of surface flow regimes of Kīlauea lava lake 71 4.3 Datasets 71 xi

Table of4.4 Contents Methods 73 4.5 Results 73 4.5.1 Passive outgassing features 73 4.5.1.1 Discrete Bubble Bursts 73 4.5.1.2 Surface Doming 74 4.5.1.3 Spot Vents 76 4.5.2 Distribution of discrete bubble bursts (boundary and intraplate) 77 4.6 Discussion 81 4.6.1 Relationships among of passive outgassing processes 81 4.6.2 Driving mechanism behind passive outgassing – Surficial or deep processes? 83 4.6.3 Implications for deeper lava lake processes 85 4.7 Conclusions 87

5. Summary and Future Resarch 88 5.1 Conduit processes and dynamics 88 5.1.1 Volatile exsolution and vesiculation processes 88 5.1.2 Magma/lava interaction with external water 90 5.2 Eruption unsteadiness 91 5.3 Comparison to global examples and future research 91

6. Bibliography 94 List of Tables Table 2.1 Eruption statistics from Heliker et al. (2003) for episodes 32, 37, 40, 44 and 45 of Puʻu ʻŌʻō eruption. 35 Table 2.2 Texture and vesicularity data for all analysed pyroclasts. 40 List of Figures Figure 1.1 A) A DEM map of the island of Hawaiʻi showing the five separate volcanoes that constitute its landmass. Marked in red are the four field locations studied in this project. Inset: a map of the Hawaiian Island Chain, showing where the Island of Hawaiʻi is located in reference to the other seven major islands. 21 Figure 1.2 Top: The location of the two major rift zones of Kīlauea Volcano, the South West Rift Zone (SWRZ) and the East Rift Zone (ERZ), as well as the distribution of major eruption centres (black dots). Bottom: The location of the major fault zones of Kīlauea. In both figures the dashed line represents the contact between Mauna Loa and Kīlauea. From (Swanson et al., 1976). 22 Figure 1.3 A schematic model of the magmatic plumbing system of Kīlauea, from (Poland et al., 2013). A) Cross section through the volcano parallel to the two major rift zones. Note: horizontal scale is only approximate. B) Map view of the line B–B’ in A. 25 Figure 1.4 A schematic map showing the major geological structures of Mauna Loa Volcano. NERZ: North East Rift Zone, SWRZ - South West Rift Zone, 1: ʻAuʻau Point, 2: Puʻu Kī. 29

Figure 2.1 A) Digital Elevation Model (DEM) of Island of Hawaiʻi. B) DEM showing the summit and the approximate location of the upper part of the East Rift Zone of Kīlauea Volca- no. 33 Figure 2.2 Photographs of high Hawaiian fountains at Puʻu ʻŌʻō. 1) Incandescent, molten jet; 2) Black ash and lapilli in the cooler outer region of the jet. A) Episode 45; view to the south. B) Episode 44; view to the southwest. Note: fountain heights are peak fountain height of the episode, not fountain heights at the time the photograph is taken. Photo- graph credit: Hawaiian Volcano Observatory. 35 Figure 2.3 A) Stratigraphic column of the deposit sampled, showing eruption episode number and sample set number for each unit. B) The location of the sample site relative to the Puʻu ʻŌʻō vent. The photo was taken after episode 28 on December 15, 1984. Photo courtesy of HVO. 36 Figure 2.4 Density histograms for the pyroclast sample sets from each episode. Small numbers within the histograms correspond to the specific pyroclasts (listed on right) chosen for microtextural analysis. For comparison, blue lines are at fixed densities of 400, 800 and 1000 kg.m-3. 38 Figure 2.5 A) Binary image showing bubbly texture. B) Binary image showing foamy texture. C) Example of a pyroclast that exhibits both bubble and foamy texture. Note that that the transition between the two textural domains is sharp. D) Binary image showing irregular bubbly texture. 41 Figure 2.6 Four separate 500x magnification images showing the microlite species, abundance and textures from pyroclast 04B-38. In all images cpx: clinopyroxene, P: plagioclase, mag: magnetite and ves: vesicle. See supplementary material for locations of images within the pyroclast. 42 xiv

List of figures m Figure 2.7 Plot showing vesicularity versus vesicle number density (N v) for three high Hawaiian fountaining eruptions at Kīlauea. 42 Figure 2.8 Histograms showing vesicle size distributions for representative pyroclasts from each high fountaining episode studied. Equivalent diameter is defined as the diameter of a circle with the same area as calculated for the vesicle (Adams et al., 2006). Pyroclasts from each episode are arranged from top to bottom in order of increasing vesicularity (decreasing density). In each plot, the pyroclast number is given in the top right and the Nmv (bubbles cm-3), vesicularity (%) and density (kg cm-3) are listed in the top left. The vertical blue line through the 0.16 mm equivalent diameter is for comparison. 44

m Figure 2.9 VG/VL versus N v plots for each eruption episode showing the post-fragmentation vesiculation process that is defined by the pyroclast microtextures (after Stovall et al., 2011). In all diagrams, open circles denote primary pyroclasts and the grey arrows highlight the associated trend. The images show the microtexture of the labelled pyroclasts that are representative of textures found for that episode. A) Episodes 37, 44 and 45, all of which show growth plus coalescence trends. B) Episode 40, which shows both a growth plus coalescence trend and a growth plus nucleation trend. C) Episode 32, which has a growth trend. D) Key showing the possible trends and their associated post fragmentation vesicle processes. L = Loss, C = Coalescence, G+C = Growth plus Coalescence, G = Growth, G+N = Growth plus Nucleation, N = Nucleation. 46

m Figure 2.10 Plot of average N v versus fountain height for each eruptive episode. Black bars indi- m cate the maximum and minimum N v. Numbers correspond to the episode. 47 m Figure 2.11 A) Plot of primary pyroclast N v versus fountain height for each eruptive episode for which primary pyroclasts were identified. B) Qualitative microtexture of primary pyroclasts depicted in A). 48

Figure 2.12 Plot showing Δ(VG/VL) vs peak fountain height for all fountaining episodes sampled in this study, as well as three episodes (1, 15 and 16) from the 1969 Kīlauea Iki eruption from Stovall et al. (2011) and Stovall et al. (2012). Linear regression is calculated on all data points. 50

Figure 3.1 A: Digital elevation map of Puʻu Kī. The location of the outcrop shown in Figure 2 is indicated. On both maps, dotted lines represent the spatial extent of the cone structure and the inset images show the location of the cone on Hawaiʻi. Note: Blue corresponds to sea level. B: Digital elevation map of ʻAuʻau Point. Solid line represents escarpment encrusted with welded lapilli and bombs. The location of the outcrop shown in Figure 3 is indicated. 54 Figure 3.2 Puʻu Kī South Cone (location given in Figure 1A). Unit A: Keliuli Cone, Unit B: Inter- bedded clastogenic lavas (grey) and oxidised welded lapilli and bombs (red), Unit C: Unconsolidated ash and lapilli derived from Puʻu Kī North Cone. See text for detailed deposit descriptions. Sea cliff is approximately 20 m high. 58 Figure 3.3 Clastogenic lava (grey) and thermally oxidised, welded lapilli and bombs from Puʻu Kī South Cone. Note that there is a gradation in welding away from the clastogenic lava. CL; Clastogenic lava, DW: Densely welded lapilli and bombs, W: Welded lapilli and bombs. 59 Figure 3.4 A: A Remnant clast (indicated by arrow) within the coherent centre of a clastogenic lava lithofacies within Unit B at Puʻu Kī South Cone, also picture in Figure 3.3. B: Colour xv

List of figuresvariations within the same lithofacies as A, which are defined by the thermal oxidation of remnant clasts. 59 Figure 3.5 A) An example of a cored bomb from Puʻu Kī South Cone. The core consists of an angular, vesicular basalt clast which is coated with a rim of dense, non-vesicular, thermally oxidised basalt. The core and rim appear to differ only in their vesicularity and level of thermal oxidation. B) A 4-5 m thick layer of tack- to moderately-welded, thermally oxidised lapilli and bombs from at Puʻu Kī South Cone. The visible, well-rounded bombs are predominantly cored bombs, similar to that pictured in A. 60 Figure 3.6 ʻAuʻau Point cone (location given in Figure 1B). Unit D: Interbedded clastogenic lavas (grey) and variably welded, oxidised lapilli and bombs (red), Unit E: non-welded ash- rich unit, Unit F: layer of coarse bombs. See text for detailed deposit descriptions. Sea cliff is approximately 30 m high. FZ = fault zone. 60 Figure 3.7 Schematic of the formation of ʻAuʻau Point and Puʻu Kī South Cone. A) lava flows over saturated substrate. B: Tensile cracks form in the basal crust allowing water to percolate into the lava tube. 1. Surface crust 2. Water saturated substrate 3. Tensile fractures. C) Extensive mixing occurs between the sea-water and the molten core of the lava tube, initiating littoral lava fountaining and resulting in the deposition of clastogenic lavas. 4. Littoral lava fountaining activity 5. Clastogenic lavas 6. Steady influx of sea-water into the lava tube. D: The activity transitions into low energy lava bubble bursting activity. Heavy clast recycling within the vent results in lithofacies dominated by cored bombs. 7. Lava bubble bursting activity 8. Oxidised welded lapilli and bombs 9. Water level depletion, and reduction in water influx. 64 Figure 4.1 A: Digital Elevation Map (DEM) of Hawaiʻi. B: DEM of the summit region of Kīlauea Volcano, including the locations of the summit caldera, Halemaʻumaʻu, Overlook Vent and the East Rift Zone (ERZ). C: A close up of the lava lake within the Overlook Vent, which itself is located within Halemaʻumaʻu (image taken 9th December 2015). The dimensions of the lake are approximately 180 by 250 m. The dominant surface flow regime of the lava lake is from the NW Upwelling Zone (1) to the SW Sink (2) where spattering can be seen on the lake surface. 69 Figure 4.2 A: A far field view of the lava lake within the Overlook Vent showing the locations of the NW Upwelling Zone (NW UZ), the SW Sink as well as the general flow direction of the atable flow regime (arrow) as defined by (Patrick et al., 2016a). B: Infrared thermal images of stable flow regime (left) and unstable flow regime (right) from the HVO thermal camera. Arrow designates flow directions. C: Snapshots from high-speed infrared videos showing stable flow regime (left) and unstable flow regime (right). 72 Figure 4.3 Two separate time series, A-E and F-J showing the evolution of two examples boundary bubble bursts. In all cases the bubble bursts shown are metre scale and occur over 5-6 seconds. In each time series, the dashed boxes indicate static reference frames for each bubble burst. 75 Figure 4.4 A sequence of frames showing a 4-5 m intra-plate bubble bursting event. White dotted lined indicates the approximate extent of the bubble. 76 Figure 4.5 Frames 1 to 6 show six separate intraplate bubble bursts (~1m in size) in a fixed frame of reference on the lava lake surface. The frames span a time period of approximately three minutes. In each case the dotted line shows the approximate radius of the gas slug that impacts on the lake surface (inferred from previous frames). In each instance, the surface expression of the bubble burst and amount of incandescent lava revealed varies xvi

List of figuresgreatly despite being produced by gas slugs of similar size. 77 Figure 4.6 An example of a surface doming event. Frame 1 shows the lava lake before event. Frame 2 shows the moment at which the slug (denoted by white arrow) impacts the lava lake crust, but does not break through. Frame 3 is the same moment in time as Frame 2, but has the extent of the surface doming event highlighted by white dashed lines. 78 Figure 4.7 Bubble burst distributions for four separate highspeed videos taken of the lava lake on December 3rd 2016. Zones 1 to 3 are areas common to videos A, B and C which have the highest density of bubble bursts observed. These zones are not defined in video D asa it was much shorter in length, not allowing for enough bubble bursts to be captured. In all cases, annotation UZ denotes the NW Upwelling Zone. Note: temperature ranges are relative temperature. Definitive temperature measurements were not possible due to the masking effects of the ever-changing gas plume. 79 Figure 4.8 A & B show two separate relative temperature profiles for Zone 1 from two videos, A and B in Figure 6. These plots show The Maximum (red) Average (grey) and Minimum (blue) relative temperatures for Zone 1. Inset images 1 & 2 show typical temperature profiles. Periods of elevated maximum temperature occur as a plate boundary passes through the reference frame, punctuated by high-frequency peaks, indicating numerous boundary bubble bursts (red triangles). Periods of low maximum temperature occur when the interior of a plate occupies the reference frame. Low frequency peaks (green circles) indicate isolated intra-plate bubble bursts. The amplitude of the intra-plate bubble burst peak is proportional to the degree of crust disturbance caused by the bubble (Figure 5). 80 Figure 4.9 Bubble burst distribution for a 24hr time-lapse video of the lava lake surface. Annota- tion UZ denotes the NW Upwelling Zone. 81 Figure 4.10 Steps A to E show to formation and possible surface expressions of passive outgassing features of the lava lake at Kīlauea. Gas slugs rise through the lava within the lake (A) and impact on the viscoelastic crust at the lake surface (B) causing surface doming. If the slug does not have enough energy to break the crust, the crust rebounds (C & D) resulting in a surface doming event. If the slug possesses enough energy to break through the crust, a bubble burst occurs (E). If the crust is undisturbed, but fractures (F) gas is emitted and deflation occurs, resulting in a spot vent as the bubble moves along with surface plate motion (G). 82 Figure 4.11 A: A schematic cross-section of the Kīlauea lava lake, showing the foamy layer below the viscoelastic crust and the assumed flow with the lake (arrows), from Patrick et. al (2016a). B: Schematic showing a possible intra-lava lake source for the slugs that drive passive outgassing at the lava lake surface. In this case the peripheries of the upwelling plume interact with the foamy layer causing either localised coalescence of bubbles or a preferential pathway for bubbles rise in a ring geometry. C: A schematic showing an upper volcanic conduit source for gas slugs that drive passive outgassing at the lava lake surface. In this case bi-directional flow at the opening of the conduit at the base of the lava lake causes shear induced coalescence of bubbles at the interface between rising bubbly magma and dense, degassed magma (in a bi-directional flow regime such as core-annular flow) Meter scale slugs then buoyantly ascend trough the lake decoupled from the lake convection. In both A and B grey dotted lines represent the approximate extent of upwelling magma, arrows represent flow direction. 84

Chapter 1 Introduction

Hawaiʻi is viewed as the “type example” for many fundamental basaltic volcanic phenomena (Garcia, 2017). In the last 2500 years, activity at Kīlauea Volcano has been dominated (~60%) by periods of high frequency explosive eruptions, however the magma supply during these periods has been only 1-2% of that during periods dominated by effusive activity (Swanson et al., 2014). As a result, processes that occur during effusive periods, i.e. Hawaiian fountaining, lava lake activity and littoral rootless cone eruptions are common and typify the long-term behaviour of the volcano. Products of such activity are rarely preserved in the rock record (Moore and Ault, 1965) or are rapidly buried by products of later eruptions (Heliker et al., 2003a). Therefore, the observed historical activity at both Kīlauea and Mauna Loa provides the opportunity for the characterisation of the fundamental dynamic processes that occur during the ephemeral eruptive activity common during effusive periods. A more detailed understanding of the variables that govern the eruption styles of volcanoes can be obtained through the combination of multidisciplinary datasets and real time observations. Bonadonna et al. (2016) categorized these variables to include four broad categories: initial conditions, conduit magma dynamics, eruptive processes and parameters, and external factors. Examples of these variables include (but are not limited to) pre-eruptive volatile content (Johnson et al., 1994; Papale, 2005), the depth of degassing (Edmonds and Gerlach, 2007; Spilliaert et al., 2006), bulk magma rheology (Llewellin and Manga, 2005; Manga et al., 1998) and the rates of bubble nucleation, growth, coalescence and loss (Adams et al., 2006; Cashman and Mangan, 1994; Lautze and Houghton, 2007; Sable et al., 2006), and interaction with external water (Barberi et al., 1992; Cashman and Sparks, 2013; Koyaguchi and Woods, 1996; White et al., 2003). At Kīlauea and Mauna Loa, an extensive geophysical and physical monitoring network and both historical and current observations made by the staff at the Hawaiian Volcano Obser- 18 vatory1.1 Project (HVO) Aims have allowed for numerous basaltic eruption styles to be described in great detail by their deposit features as well as by concurrent visual observations and geophysical monitoring during the eruption. Examples include the 1983-present eruption of Puʻu ʻŌʻō (Heliker et al., 2003b; Wolfe et al., 1987), the 2010-2018 stable lava lake at Halemaʻumaʻu (Patrick et al., 2016a; Patrick et al., 2016b; Patrick et al., 2016c) and the 2008 littoral rootless cone activity at Waikupanaha (Orr, 2008). Multidisciplinary data makes it possible to study the controls on many fundamental low intensity basaltic eruption styles in Hawaiʻi (Poland, 2017). These include Hawaiian style fountaining, persistent lava lake activity and rootless littoral cone activity. Datasets such as these are vital in quantifying the processes that drive these eruptions in order to better define fundamental eruption styles that constitute the full spectrum of activity typical of basaltic volcanoes, thus, reducing the difficulties in quantitatively comparing a single style of eruption across two separate locations

1.1 PROJECT AIMS

The aim of this project is to identify and quantify the processes and parameters that can influence the eruption style of basaltic volcanism in the uppermost stages of the magma ascent pathway, as well as the surface transport of lava within lava-tubes, to provide a better understanding of the full spectrum of activity common to basaltic volcanoes. This study presents three case studies focussed on separate examples of ephemeral basaltic eruptive activity in Hawaiʻi: five episodes of Hawaiian-style fountaining in the 1983-present eruption of Puʻu ʻŌʻō on Kīlauea, the 2010-2018 stable Halemaʻumaʻu lava lake at the summit of Kīlauea, and the rootless littoral cones at Puʻu Kī and ʻAuʻau Point on Mauna Loa, each of which is focussed on a different phenomenon.

1. Hawaiian-style Fountaining This study compares five of the later high Hawaiian fountaining episodes from Puʻu ʻŌʻō in 1985-1986, which erupted magma of near constant composition but had varying fountain heights and durations. The vesiculation histories through microtextural analysis of tephra erupted during episodes 32, 37, 40, 44 and 45 are quantified.The microtextural characteristics of the deposits are then linked to visual observations of fountain geometries made by the staff of the U.S. Geological Survey’s Hawaiian Volcano Observatory (HVO). The data are used to answer three key questions: (1) do lapilli-sized pyroclasts 19

1.1 Projectfrom these Aims episodes retain any signature of syn-eruption magma properties? (2) does fountain height (reflecting decompression rate) scale with bubble number density of lapilli? and (3) is there any quantifiable textural signature (such as vesicle-to-melt ratio,

VG/VL) that can be used as a proxy for clast residence times, and therefore fountain height, for Puʻu ʻŌʻō fountains?

2. Rootless littoral cone eruptions The aim of this study is to critically assess confined mixing as a model of genesis for the rootless littoral cones at Puʻu Kī and ʻAuʻau Point in order to build upon the deposit description and model of formation put forward by Jurado-Chichay et al. (1996a). We also consider the similarities and differences between Puʻu Kī South Cone and ʻAuʻau Point and Icelandic rootless cones to determine the implications these two cones have on criteria for discriminating between ‘dry’ deposits and the deposits of lava–water interactions.

3. Stable lava lake activity This study utilises high-speed thermal infra-red videos and thermal image time-lapses of the surface of Kīlauea lava lake to document the small-scale, transient degassing features such as bubble bursts and spot vents. By mapping the spatial distribution and temporal frequency of these features, we aim to answer two key questions: 1) are these features governed by shallow or deep-seated degassing processes? and 2) can these small-scale features be used to better understand the large-scale processes, such as magma ascent, degassing and convection, within the lava lake and/or shallow conduit of Kīlauea volcano?

The results of each of these three individual studies can then be used as a framework to better understand two key phenomena; conduit processes and dynamics and eruption unsteadiness. A better understanding of both of these phenomena is a critical building block in working towards developing a general, descriptive classification scheme of low intensity basaltic volcanism (Bonadonna et al., 2016) as well as giving us the ability to compare the complexities of these eruption styles across separate locations around the world. 20

1.2 Volcanological VOLCANOLOGICAL Settings SETTING The eight primary Hawaiian Islands constitute the southernmost extent of a far greater volcanic chain known as the Hawaiian – Emperor Seamount Chain (HESC) (Clague and Dalrymple, 1987; Langenheim and Clague, 1987). The Hawaiian – Emperor Seamount Chain stretches approximately 6000 km across the northern Pacific Ocean and consists of at least 129 shield volcanoes (Clague and Dalrymple, 1987; Sharp and Clague, 2006). In total, the volume of the HESC is believed to be greater than 1×106 km3, although beyond Kure Island, none of the Seamounts in the HESC protrude above sea level (Bargar and Jackson, 1974; Clague and Dalrymple, 1987). The island of Hawaiʻi is the largest and southernmost island in the Hawaiian Island chain. Hawaiʻi is made up of at least seven different, coalesced volcanoes Fig.( 1.1), five of which are above sea level (Kohala, Mauna Kea, Hualālai, Mauna Loa, Kīlauea) and two that are submarine seamounts (Lōʻihi and Mahukona) (Moore and Clague, 1992; Morgan et al., 2010). Some of these volcanoes formed in isolation, but the majority of the volcanoes that constitute the island of Hawaiʻi grew on the flanks of, or in close proximity to, existing volcanoes. This is the result of the movement of the Pacific Plate over the Hawaiian hot-spot, shifting the loci of new volcano formation progressively to the south. The net result of this is that the Island of Hawaiʻi consists of volcanoes that are coalesced and sometimes interleaved stratigraphically if adjacent volcanoes were active contemporaneously (Holcomb et al., 2000; Park et al., 2009).

1.3 THE GEOLOGY OF KĪLAUEA VOLCANO

1.3.1 Structure of Kīlauea Volcano Kīlauea volcano is located on the south eastern shore of the island of Hawaiʻi (Fig. 1.1), covering approximately 1500 km2 and has a summit elevation of approximately 1.24 km above sea level (Rowan and Clayton, 1993). Kīlauea is split into northern and southern flanks by two rift zones that emanate from the summit caldera (Fig. 1.2), the Southwest Rift Zone (SWRZ) and the East Rift Zone (ERZ) (Klein et al., 1987). These rift zones are visible in the field as topographic ridges and are the loci for the majority of activity on the volcano which is not erupted from the summit (Borgia, 1994; Rowan and Clayton, 1993).The northern flank and the SWRZ are generally less active and more stable, when compared to the southern flank and the ERZ, due to the increased effect of Kīlauea onlapping Mauna Loa along its northwestern extent 21

1.3 The Geology of Kīlauea Volcano

Figure 1.1 – A) A DEM map of the island of Hawaiʻi showing the five separate volcanoes that constitute its landmass. Marked in red are the four field locations studied in this project. Inset: a map of the Hawaiian Island Chain, showing where the Island of Hawaiʻi is located in reference to the other seven major islands.

(Rowan and Clayton, 1993). As well as the two rift zones, the other prominent surface structures of Kīlauea are the three major fault systems that are active on the volcano. These fault systems are the Koae Fault System, the Hilina Fault System and the Kaoiki Fault System (Fig. 1.2). The Koae system is 12 km long and 2 km wide (Duffield, 1975), located directly south of the summit caldera. It is comprised of predominantly vertical dipping fractures which strike ~75°, and intersects the two rift zones (Duffield, 1975; Rowan and Clayton, 1993). The Koae system is believed to be a passive tear away zone (Duffield, 1975), that along with the rift zones, accommodates the southern migration of the southern flank of Kīlauea due to the forceful intrusion of magma into the rift 22

1.3 The Geology of Kīlauea Volcano between Mauna Loa and Kīlauea. From (Swanson et al., 1976). A) The location of the two major rift zones of Kīlauea Volcano, the South West Rift Zone (SWRZ) and the East Rift Zone the West (ERZ), South as well as the The A) location of the two major rift zones Volcano, of Kīlauea – distribution of major The eruption location centres of (black the dots). major B) fault zones of Kīlauea. In both figures the dashed line represents the contact Figure 1.2 Figure 23 zones1.3 The from Geology the summit of Kīlauea magma Volcano storage zone (Swanson et al., 1976). The Hilina Fault System is located on the southern, seaward flank of Kīlauea with its northern extent clearly defined by the large pali (Hawaiian for steep hill or cliff). It is comprised of a series of normal faults, dipping to the southwest, caused by the gravitation slip of the seaward edge of Kīlauea (Crosson and Endo, 1982; Swanson et al., 1976). The Kaoiki Fault Sys- tem is parallel to the SWRZ, and is predominantly located on the southwestern flank of Mauna Loa but is also active on Kīlauea’s northern flank with such features as the Kaoiki Pali (Koyan- agi et al., 1966; Rowan and Clayton, 1993). In recent history, there has been no record of any major slip along the Kaoiki Pali , however, numerous studies have recorded regular movement along the Hilina Pali including the Kalapana Earthquake in 1975 and a 6.9 Mw earthquake on the 4th May 2018 (Brooks et al., 2006; Denlinger and Okubo, 1995; Lipman et al., 1985; Liu et al., 2018; Wolfe et al., 2007). The Koae Fault System has been constantly slipping since 1970, resulting in major southward migration of the southern flank of Kīlauea (Delaney et al., 1993; Denlinger and Okubo, 1995). The extent of the combined deformation along the rift zones and the southern flank of Kīlauea is larger than expected to be caused solely by shallow dyke injection from a summit magma reservoir (Delaney et al., 1990). It is also too large to be solely due to gravitational forces, and thus, is rather a combination of gravity and forceful intrusion of magma into the rift zones (Dvorak et al., 1986; Swanson et al., 1976). This led to the hypothesis that there is a deep flank instability, or decollement, accommodating the southern flank movement to the southeast (Borgia, 1994; Denlinger and Okubo, 1995; Got et al., 1994). This decollement becomes evident in datasets that document the focal point of earthquakes at Kīlauea over a large period of time. Such datasets show that there are two distinct zones of seismicity with one at around 1.5 km depth and another at around 7 km, depicting the shallow and deep regions of the magma plumbing system. There is also a sudden drop off in the occurrence of earthquakes on the southern flank of Kīlauea below a depth of 10-12 km, believed to be caused by the deep decollement (Denlinger and Okubo, 1995; Got et al., 1994; Klein et al., 1987).

1.3.2 The Magmatic Storage System of Kīlauea Volcano Figure 1.3 outlines a contemporary model for the magmatic plumbing system that underlies Kīlauea, formulated and presented by Poland et al. (2013). This model highlights the 24 existence1.3 The Geology of three of primary Kīlauea summit Volcano reservoirs: the Halema’uma’u Source, the South Caldera Source and the Keanakāko’i Source. Not only are there shallow magma reservoirs under both the SWRZ and the ERZ, but there exists another, smaller reservoir below Kīlauea Iki called the Halema’uma’u-Kīlauea Iki Trend and a deeper, less volcanically active reservoir below the SWRZ called the Seismic SWRZ.

Kīlauea Summit Reservoirs Geochemical modelling has been used to estimate the geometry of the magma storage zone beneath the summit of Kīlauea. These models suggest that the volume of melt below the caldera is approximately 2-3 km3, which is much less than any geophysical estimate of 2-40 km3 (Poland et al., 2013). This is because geophysical models are not able to distinguish between melt, crystal mush, and the surrounding hydrothermal system (Pietruszka and Garcia, 1999). The shape of the summit storage is much better described by modern seismic and microgravity measurements. The primary storage zone beneath Kīlauea’s summit, the South Caldera source (SC, Fig. 1.3) is fed by a near-vertical conduit that stretches to approximately 15 km below the surface (Wright and Klein, 2006). The magma source is between 3 km (Poland et al., 2013) and 7 km (Wright and Klein, 2006) deep and is the initial storage zone for all magma that is erupted at Kīlauea. There are also two smaller summit magma sources; the Keanakāko’i and Halema’uma’u sources (K and H in Fig. 1.3 respectively). The Keanakāko’i source is located approximately 2.6 km directly below the Keanakāko’i Crater (Poland et al., 2013) and the Halema’uma’u Source is between 1-1.4 km below the northeastern edge of the Halema’uma’u Crater (Almendros et al., 2002; Poland et al., 2013).

Rift Zone Magma Sources Both major rift zones at Kīlauea are underlain by at least one active magma source in the form of a shallow dyke. The ERZ has been the loci of most modern rift zone activity (such as the ongoing Puʻu ʻŌʻō-Kūpaianaha eruption) and its subaerial expression stretches from the Kīlauea caldera to the eastern tip of the Island, beyond Kapoho Crater. Geophysical modelling suggests that the magma source of the ERZ lies at approximately 2-3 km below the surface, links up with the Halema’uma’u summit reservoir, and has a molten core that stretches its entire length (Fig. 1.3) (Poland et al., 2013). Physical observations made during eruption episodes at 25

1.3 The Geology of Kīlauea Volcano view of the line B–B’ in A. in view of the line B–B’ A schematic model of the magmatic plumbing model of the magmatic schematic A – Figure 1.3 Figure system of Kīlauea, from (Poland et al., 2013). A) Cross rift two major the to parallel volcano the through section Map B) approximate. only is scale horizontal Note: zones. 26

Puʻu1.3 The ʻŌʻō Geology support of a Kīlauea molten conduitVolcano between the summit and the ERZ. Inflation-deflation measurements taken simultaneously at Puʻu ʻŌʻō and the summit during early fountaining episodes in the early 1980s suggest that the magma being erupted was being fed directly to Puʻu ʻŌʻō from the summit reservoirs at the time of eruption, confirming the existence of an open conduit below the ERZ (Montagna and Gonnermann, 2013). Modern fissure eruptions, such as the 2011 Kamoamoa eruption, are thought to be the product of dykes branching from the ERZ magma source and intruding through the final kilometres and intersecting the surface (Lundgren et al., 2013; Orr, 2013), whereas Puʻu ʻŌʻō is linked to the ERZ magma source via a sustained conduit during the last 30 years (Poland et al., 2013). As well as a continuous magma source at 2-3 km depth, there are isolated pockets of magma remaining from past dyke intrusions into the upper regions of the rift zone (Fig. 1.3) (Poland et al., 2013). Unlike the ERZ, the SWRZ is underlain by two separate magma bodies, the Volcanic SWRZ and the seismic SWRZ. The volcanic SWRZ is at a depth of ~1 km extending out from the Halema’uma’u summit reservoir and is the source of many historical eruptions such as Mauna Iki (Poland et al., 2013). The Seismic SWRZ lies at a much greater depth (2-3 km), is linked to the South Caldera Source, and is marked by a zone of high seismic activity related to dyke intrusions and movement along the main boundary between the stable north and mobile south flanks of the volcano (Poland et al., 2013). The final magma reservoir is the shallow plumbing of Kīlauea is the Halema’uma’u-Kīlauea Iki Trend (Fig. 1.3). Although it is not part of either rift zone, this magma body is manifested on the surface as a number of northeast striking fissures that run away from Halema’uma’u on the summit caldera floor (Zurek and Wil- liams-Jones, 2013). The magma body is around 1 km deep and has been the source of numerous modern eruptions including the 1959 Kīlauea Iki fountaining episodes (Stovall et al., 2012)

1.3.3 1983-2018 Activity at Kīlauea Volcano The Puʻu ʻŌʻō eruption began on the 3rd of January 1983. After three months dominated by distributed fissure fountaining, activity localised at a vent later named Puʻu ʻŌʻō (Heliker et al., 2003b; Wolfe et al., 1988). From June 1983 to June 1986, 44 Hawaiian style fountaining episodes took place at Puʻu ʻŌʻō. The activity was highly cyclic during this time period, with lengthy repose periods (8 to 65 days) which were punctuated by relatively short (0.3 to 16 days) Hawaiian fountaining episodes. Repose periods were characterised by fluctuating magma levels within the conduit and outgassing from the vent during gradual inflation of Kīlauea’s sum- 27 mit.1.3 The Eruptive Geology episodes of Kīlauea consisted Volcano of low (≤200 m) to high (>200 m) Hawaiian fountaining and lava effusion (both pāhoehoe and ʻaʻā lavas), accompanied by rapid deflation at the summit of Kīlauea (Heliker et al., 2003b; Wolfe et al., 1987; Wolfe et al., 1988). This pattern was broken in July 1986, when the eruption shifted to a new vent (Kupaianaha) and the eruptive style changed to one of nearly continuous effusion (Heliker et al., 2003b). Orr et al. (2015) outline in detail the ongoing eruption of Puʻu ʻŌʻō from June 1986 to present day. This period includes a further thirteen episodes overwhelmingly dominated by continuous effusion punctuated by changes in vent location along the ERZ in the immediate vicinity of Puʻu ʻŌʻō, repose periods of varying length, and fissure eruptions (see table 18.2 of Orr et al. 2015). On March 19th 2008 summit activity resumed at Kīlauea, simultaneously with the eruption at Puʻu ʻŌʻō in the ERZ. The recent period of sustained lava lake activity at Kīlauea began on March 19th 2008 when a small explosion occurred simultaneously with the opening of a 35 metre vent along the eastern rim of Halemaʻumaʻu (Wilson et al., 2008). This explosion was preceded by several months of increased seismic activity and SO2 emissions (peaking at >1600 tonnes.day-1) (Houghton et al., 2011; Wilson et al., 2008). Lava first became visible within the ‘Overlook Vent’ in 2008 and continued between 2008 and 2009 as ephemeral lava ponds (Orr et al., 2013; Patrick et al., 2016b). By February 2010 a larger, continuously active lava lake had appeared (Patrick et al., 2016b) and by 2011 had grown into a 150 m wide lava lake that extended into Halemaʻumaʻu (Orr et al., 2013). Over the next five years the lava lake grew to its maximum dimensions of approximately 180x250m and had surface levels routinely between 30-60m below the crater rim (Patrick et al., 2016a). On the 1st of May 2018 deflation of the summit of Kīlauea began, followed by the initiation of the lava lake level dropping within Halemaʻumaʻu (Hawaiian Volcano Observatory, 2018). On the 3rd of May 2018 a new eruptive phase began in the lower ERZ of Kīlauea, whilst the lava lake surface within Halemaʻumaʻu continued to drop due to the migration of magma from the summit reservoirs to the ERZ (Ha- waiian Volcano Observatory, 2018; Liu et al., 2018). The lava lake continued to drop until the 10th of May 2018 when the Halemaʻumaʻu lava lake disappeared from view (Hawaiian Vol- cano Observatory, 2018). The eruption on the lower ERZ of Kīlauea continued into late 2018 until activity ceased in September (Hawaiian Volcano Observatory, 2018; Liu et al., 2018; Neal et al., 2019). 28

1.4 The THE Geology GEOLOGY of Mauna OF MAUNA Loa Volcano LOA VOLCANO

1.4.1 Structure of Mauna Loa Volcano

Mauna Loa Volcano is the largest volcano on earth, with an estimated volume of 40,000 km3 (Lockwood and Lipman, 1987; MacDonald, 1972). At its summit, Mauna Loa has a collapse caldera known as Mokuʻaweoweo that is approximately 3 km by 5 km (Fig. 1.4) (Lock- wood and Lipman, 1987). Two rift zones extend from Mokuʻaweoweo (Fig. 1.4), the Southwest Rift Zone (SWRZ) and the Northeast Rift Zone (NERZ) (Lipman, 1980a; Lipman, 1980b; Lockwood and Lipman, 1987). The NERZ is approximately 50 km long, however it likely ex- tends below the flank of Kīlauea Volcano, with historical activity being localised to the upper 19 km (Lockwood and Lipman, 1987). The SWRZ has a subaerial extent of 65 km and continues a further 35km below sea level to the south (Lockwood and Lipman, 1987). As well as these two rift zones, eruptions have occurred within a zone of radially distributed vents to the NW of Mokuʻaweoweo. Historic eruptions are either classified as summit or flank eruptions (although a single eruption can have a summit phase and a flank phase) with 31% occurring in the NERZ, 25% in the SWRZ, 6% at the NW radial vents and 38% at the summit region (Lockwood and Lipman, 1987; Walter and Amelung, 2006). Mauna Loa has a magma reservoir below its summit at a depth of approximately 3.5-5 km (Decker et al., 1983; Miklius et al., 1995; Okubo et al., 1997; Rhodes, 1988; Walter and Amelung, 2006). As with Kīlauea Volcano, eruptions at Mauna Loa are initiated from the summit reservoir, and magma is intruded along the rift zones and erupted on its flanks (Lockwood and Lipman, 1987; Rhodes, 1988; Walter and Amelung, 2006). Dykes that intrude down the rift zones do so at an average depth of 3-5km (Walter and Amelung, 2006).

1.4.2 Geology and Historic Eruptions of the Southwest Rift Zone of Mauna Loa

The geology of the SWRZ of Mauna Loa is split into four main units: The Kaʻu Volca- nics, the Kahuku Volcanics, the Ninole Basalts and the Pahala Ash (Lipman, 1980a; Lipman et al., 1990; Lipman and Swenson, 1984; Stearns et al., 1930). The Kaʻu Volcanics span 30 ka to present, including the historic rock units from the 19th and 20th centuries (Lipman and Swenson, 1984). The pre-historic sequences of the Kaʻu Volcanics are split into five separate units by Lipman (1980a) based on stratigraphic sequence. The rocks featured in this study, at AuʻAuʻ Point and Puʻu Kī, belong to age-group three ranging from 750-1500ya. 29

1.4 The Geology of Mauna Loa Volcano Figure 1.4 – A schematic map showing the major geological structures of Mauna Loa Volcano. NERZ: North East Rift Zone, SWRZ - South West Rift Zone, 1: ʻAuʻau Point, 2: Puʻu Kī. Adapted from Neal and Lockwood (2003) and Lipman et al., (1990).

The Pahala Ash is a weathered basaltic ash unit, up to 10 m thick, that separates the younger Kaʻu Volcanics and Kahuku Volcanics. The Pahala ash was deposited at least 31,100 ± 900ya (Lipman and Swenson, 1984). The Kahuku Volcanics are similar in composition and type to the Kaʻu Volcanics, however they underlie the Pahala ash and are therefore >30ka. No source vent for the Kahuku Volcanics has ever been identified (Lipman and Swenson, 1984). The Ninole Basalts are some of the oldest rocks found to outcrop on Hawaiʻi, similar to some of those from Kohala Volcano, and are dated at 0.1-0.2 Ma (Lipman et al., 1990; Lipman and 30

Swenson,1.4 The Geology 1984). of Mauna Loa Volcano Historically, Mauna Loa has erupted 33 times (since observations were first made in 1841) with a large eruption (>1.0×108 m3) every ten years (Walter and Amelung, 2006). Of those 33, seven have occurred on the SWRZ with an average duration of 16 days and a cumula- tive volume of 8.6×105m3 (Lockwood and Lipman, 1987), with flows often reaching the ocean (Moore and Ault, 1965). Eruptions have occurred along almost the entirety of the SWRZ, with the lowest elevation historic eruption occurring at 640m above sea level (Lipman, 1980a). Ap- proximately half way down its subaerial length, the SWRZ has a broad 40° bend. This bend has been the source of many of the most voluminous eruptions in the SWRZ (the greatest of which occurred in 1950) forming a large shield around Puʻu o Keokeo (Lipman, 1980b). Chapter 2 Eruption and fountaining dynamics of selected 1985–1986 high fountaining i. episodes at Kīlauea volcano, Hawaiʻ

Authors

S.J. Holt, R.J. Carey, B.F. Houghton, T. Orr, J. McPhie, S. Feig

Published in Journal of Volcanology and Geothermal Research 369 (2019): 21-34. This research was supported by USGS grant SV-ARRA-0004, ARC grants DP110102196 and DE150101190. All SEM imaging was carried out at the Central Science Laboratory at the University of Tasmania.

2.1 INTRODUCTION

The dynamic processes within basaltic magma during its ascent in the shallow conduit of a volcano have a major influence on subsequent eruptive behaviour (Gonnermann and Man- ga, 2013; Houghton et al., 2004a; Houghton et al., 1999; Papale, 1998; Parfitt, 2004; Vergniolle and Jaupart, 1990). Microtextural analysis of pyroclasts is commonly used to quantify rates and timing of vesiculation processes and shallow conduit dynamics during eruptions (Adams et al., 2006; Carey et al., 2012; Costantini et al., 2010; Lautze and Houghton, 2007; Parcheta et al., 2013; Porritt et al., 2012; Rust and Cashman, 2011; Sable et al., 2006; Stovall et al., 2011; Stovall et al., 2012). Microtextural datasets can be used to quantify processes during a single episode or explosion (Carey et al., 2012; Sable et al., 2006) and also to describe the range of eruptive behaviour of a volcano during multiple eruption episodes over an extended period of time (Stovall et al., 2011; Stovall et al., 2012). Hawaiian fountaining is defined as the sustained jetting of gas and fluidal coarse pyroclasts (Taddeucci et al., 2015) and is a common phenomenon at basaltic volcanoes (Fedotov et al., 1980; Heliker et al., 2003; Polacci et al., 2006; Di Muro et al., 2016;). Kīlauea Volcano is the type-locality for Hawaiian fountaining (Fig. 2.1). In its written historical period (since 1823), 32

Kīlauea2.1 Introduction has had three well-documented eruptions that produced multiple Hawaiian fountaining episodes: at Kīlauea Iki in 1959 (Richter et al., 1970; Stovall et al., 2011; Stovall et al., 2012), at Mauna Ulu in 1969 (Parcheta et al., 2013; Swanson et al., 1979) and at Puʻu ʻŌʻō from 1983 to 1986 (Heliker et al., 2003; Wolfe et al., 1987). Hawaiian fountains constitute a low-energy end-member of sustained, explosive volcanism (Walker, 1973). Traditionally, the preferred model for the vesiculation style and dynamics of Hawaiian-style volcanism was based on the rapid ascent of the magma within the shallow conduit, permitting mechanical coupling of the shallowest-exsolved and smallest bubbles of magmatic volatile gases with the liquid melt (Cashman and Mangan, 1994; Parfitt and Wilson, 1995; Parfitt, 2004; Wilson, 1980; Wilson and Head, 1981). However, a lack of correlation between the mass discharge rate and eruption style, as well as observed rapid fluctuations in the eruption intensity during a single episode, show that eruptive behaviour is not just dependent on the rise rate of the magma, but also on the depth, timing, and rate of volatile exsolution, and the buoyant ascent of large vesicles within the shallow conduit (Gonnermann and Manga, 2007, 2013; Houghton and Gonnermann, 2008; Taddeucci et al., 2015). Variables including (but not limited to) the pre-eruptive volatile content (Johnson et al., 1994; Papale, 2005), the depth of degassing (Edmonds and Gerlach, 2007; Spilliaert et al., 2006), bulk rheology (Llewellin and Manga, 2005; Manga et al., 1998) and the rates of bubble nucleation, growth, coalescence and loss (Adams et al., 2006; Cashman and Mangan, 1994; Lautze and Houghton, 2007; Sable et al., 2006) are complex and interrelated, and govern the explosive eruption styles of basaltic volcanoes. It is possible through microtextural analysis of the erupted tephra to quantify aspects of the vesiculation history, i.e., the rates of bubble nucleation, growth, coalescence and escape in the magma within the shallow conduit (both pre- and post-fragmentation) (Adams et al., 2006; Cashman and Mangan, 1994; Lautze and Houghton, 2005; Mangan and Cashman, 1996; Parcheta et al., 2013; Polacci, 2005; Sable et al., 2006; Stovall et al., 2011; Stovall et al., 2012). The quantification of vesicularity has also been used to identify the locations, cooling rates and residence times of pyroclasts within Hawaiian fountains (Porritt et al., 2012). Progression of post-fragmentation expansion from fluidal microvesicular dense clasts to pumice in a single fountaining event was reported in Stovall et al. (2011, 2012). This study compares five of the later high Hawaiian fountaining episodes from Puʻu ʻŌʻō in 1985-1986, which erupted magma of nearly constant composition but had varying foun- 33

2.2 Eruption history of Puʻu ʻŌʻō

Figure 2.1 – A) Digital Elevation Model (DEM) of Island of Hawaiʻi. B) DEM showing the summit and the approximate location of the upper part of the East Rift Zone of Kīlauea Volcano. tain heights and durations. We quantify the vesiculation histories through microtextural analysis of tephra erupted during episodes 32, 37, 40, 44 and 45. The microtextural characteristics of the deposits are then linked to visual observations of fountain geometries made by the staff of the U.S. Geological Survey’s Hawaiian Volcano Observatory (HVO). The data are used to answer three key questions: (1) do lapilli-sized pyroclasts from these episodes retain any signature of syn-eruption magma properties? (2) does fountain height (reflecting decompression rate) scale with bubble number density of lapilli? and (3) is there any quantifiable textural signature (such as vesicle-to-melt ratio, VG/VL) that can be used as a proxy for clast residence times, and therefore fountain height, for Puʻu ʻŌʻō fountains?

2.2 ERUPTION HISTORY OF PUʻU ʻŌʻŌ

The Puʻu ʻŌʻō eruption began on the 3rd of January 1983. After three months dominated by distributed fissure fountaining, activity localised at a vent later named Puʻu ʻŌʻō Fig.( 2.1) (Heliker et al., 2003; Wolfe et al., 1988). From June 1983 to June 1986, 44 Hawaiian fountaining episodes took place at Puʻu ʻŌʻō. The activity was highly cyclic during this time period, with lengthy repose periods (8 to 65 days) punctuated by relatively short (0.3 to 16 days) Ha- waiian fountaining episodes. Repose periods were characterised by fluctuating magma levels within the conduit and outgassing from the vent during gradual inflation of Kīlauea’s summit. Fountaining episodes consisted of low (≤200 m) to high (>200 m) Hawaiian fountaining and 34 lava2.3 Sampling effusion and(both Analytical pāhoehoe methodsand ʻaʻā lavas), accompanied by rapid deflation at the summit of Kīlauea (Heliker et al., 2003; Wolfe et al., 1987; Wolfe et al., 1988). This pattern was broken in July 1986, when the eruption shifted to a new vent (Kupaianaha), and the eruptive style changed to one of nearly continuous effusion (Heliker et al., 2003).

2.2.1 Studied eruption episodes The five episodes discussed in this paper (episodes 32, 37, 40, 44 and 45) took place between April 1985 and May 1986. These episodes were chosen because all exceeded 200 m in height (e.g., Fig. 2.2), had products that could be precisely linked to an episode and sampled, and were documented in detail during eruption. The visual observations are summarised in Ta- ble 2.1, modified from Heliker et al. (2003). Together, these episodes erupted a total of 4.5×107 m3 (Dense Rock Equivalent, DRE) of magma and had fountain heights ranging from 264 to 391 m. In each episode, the maximum fountain height was reached after the simultaneous gradual build up in intensity and summit deflation. At the end of each episode, seismic tremors within the ERZ decreased rapidly and summit deflation switched back to inflation (Heliker et al., 2003). The geochemical composition of the magma became steadily more mafic through the five episodes studied and contained no traces of the early hybrid magma that was erupted at Puʻu ʻŌʻō until Episode 30 (Garcia et al., 1992; Moore et al., 1980).

2.3 SAMPLING AND ANALYTICAL METHODS

2.3.1 Field work and sample selection A suite of samples was collected on the 5th of March 2008 from a 56-cm-thick deposit approximately 1 km north of Puʻu ʻŌʻō that included beds corresponding to six eruptive episodes (Fig. 2.3). At the time of sample collection, this location was one of very few where parts of the 1983-1986 pyroclastic succession was accessible. However, it was overrun by lava by 2013–2016. At this location, the stratigraphy of the tephra was described, and samples of 100 clasts in a restricted size fraction (16 to 32 mm in diameter) were collected from narrow vertical intervals within each fall unit. These samples were used for componentry analysis, juvenile clast density, and more detailed microtextural work. Tephra units from most of the Puʻu ʻŌʻō episodes are not present at this site, as the dominant direction of pyroclastic fall was southwest due to the prevailing trade winds. Because of the infrequency of tephra fall north of the vent, it 35

2.3 Sampling and Analytical methods

Figure 2.2 – Photographs of high Hawaiian fountains at Puʻu ʻŌʻō. 1) Incandescent, molten jet; 2) Black ash and lapilli in the cooler outer region of the jet. A) Episode 45; view to the south. B) Episode 44; view to the southwest. Note: fountain heights are peak fountain height of the episode, not fountain heights at the time the photograph is taken. Photograph credit: Hawaiian Volcano Observatory

Table 2.1 – Eruption statistics from Heliker et al. (2003) for episodes 32, 37, 40, 44 and 45 of Puʻu ʻŌʻō eruption. Episode Episode Area Max. Repose Episode Sample Start End Covered DRE Volume3 Fountain Episode Period2 Length Number (Date – (Date – by Lava (106 m3) Height (days) (days) Time1) Time1) (km2) (m) 21/04/85 22/04/85 32 04B 38.4 0.8 4.8 11.4 391 – 1516 – 0906 24/09/85 25/09/85 37 04A 21.8 0.5 4.4 10.3 352 – 1808 – 0619 01/01/86 02/01/86 40 03 48.5 0.6 4 8.1 264 – 1309 – 0238 13/04/86 14/04/86 44 02 22.2 0.5 5.2 8.1 308 – 2054 – 0756 07/05/86 08/05/86 45 01 23.6 0.5 5.5 6.6 257 – 2241 – 1106 Data from Heliker et al. (2003) 1 Hawaii-Aleutian Time Zone, UTC/GMT -10 2 Repose period leading up to episode. 3 Dense Rock Equivalent Volume 36 was2.3 Sampling therefore andpossible Analytical to identify methods the episodes sampled with a high degree of certainty. The deposit of Episode 30 was not studied as we were unable to obtain a sample wholly representative of the deposit. Table 1 outlines the eruption characteristics of each of the five episodes sampled, modified from Heliker et al. (2003).

2.3.2 Analytical methods The bulk density of every pyroclast within each sample was measured following the method of Houghton and Wilson (1989), measurements have a ±30 kg.m-3 error, and histograms were generated (Fig. 2.4). Bulk vesicularity for each pyroclast was then calculated assuming a DRE density of 2950 kg m-3 (Ryan, 1988). Three or four pyroclasts, representing minimum, maximum and modal densities, were then selected from each of the five sample sets for microtextural image analysis (Houghton and Wilson, 1989). Thin-sections were made from the

Figure 2.3 – A) – Stratigraphic column of the deposit sampled, showing eruption episode number and sample set number for each unit. B) The location of the sample site relative to the Puʻu ʻŌʻō vent. The photo was taken after episode 28 on December 15, 1984. Photo courtesy of HVO. 37 selected2.4 Results pyroclasts and nested image sets for each pyroclast were generated using both a Dell V105 scanner and FEI Quanta 600 MLA Scanning Electron Microscope. Images were taken at 4.5 cm, 5.0 mm and 1.25 mm horizontal field widths (corresponding to 25x, 54x and 215x magnification) and nested to ensure that full ranges of vesicle sizes were effectively imaged (Shea et al., 2010). All SEM images had a resolution of 1024 by 884 pixels and scales of 47 pixels. mm-1, 202 pixels.mm-1 and 805 pixels.mm-1 respectively for the 4.5 cm, 5.0 mm and 1.25 mm horizontal field widths. At all magnifications, the cut-off for smallest object measured was 20 pixels. Each nest contained from seven to twenty-one images. Adobe Photoshop CS6 software was used to generate binary images which were processed using the ImageJ software package (Abràmoff et al., 2004) to measure the number of vesicles per unit area and their individual size and shape. These measurements were converted to three dimensions (vesicle volume density,

Nv) using stereological techniques (Sahagian and Proussevitch, 1998). A detailed description of the image processing and microtextural analysis methodology used can be found in Adams et al. (2006), Shea et al. (2010), Stovall et al. (2011), and Parcheta et al. (2013).

2.4 RESULTS

2.4.1 Juvenile clast density and vesicularity

Figure 2.4 shows the density distributions for the pyroclast sample sets. Three episodes, episode 32, 37 and 44 have positively skewed density distributions; modal densities are between 300 and 500 kg m-3, corresponding to vesicularities of 89% to 82%. Of these three episodes, the density distributions of episode 37 and episode 44 pyroclasts are unimodal, whereas the episode 32 data are bimodal and have a secondary mode at 700 to 900 kg m-3, corresponding to vesicularities of 74% to 67%. Episode 32 has a much larger range of pyroclast densities (70–1330 kg m-3) than episodes 37 and 44 (240–1170 kg m-3 and 110–970 kg m-3, respectively). Although all three sample sets are positively skewed toward low density/high vesicularity, the episode 32 pyroclasts have a much larger sub-population of pyroclasts with densities greater than 800 kg m-3 (less than 71% vesicularity). Episode 40 pyroclasts have a modal density similar to those of episodes 32, 37 and 44 (300–400 kg m-3 and 82–89% vesicularity) but a much flatter density distribution, whereby pyroclasts with densities between 200 and 700 kg m-3 (74% to 93% vesicularity) are of more equal abundance. Episode 40 pyroclasts also have the narrowest density range of the five episodes; 38

2.4 Results

Figure 2.4 – Density histograms for the pyroclast sample sets from each episode. Small numbers within the histograms correspond to the specific pyroclasts (listed on right) chosen for microtextural analysis. For comparison, blue lines are at fixed densities of 400, 800 and 1000 kg.m-3. all 100 pyroclasts having densities between 200 and 900 kg m-3 and vesicularities between 92% and 67%. Episode 45 is the only episode that produced pyroclasts with a negatively skewed density distribution. The modal density is between 800 and 900 kg m-3 (70-67% vesicularity) and less well defined than for the three strongly positively skewed episodes (episodes 32, 37, 44). In addition to having the highest modal density of all five episodes, episode 45 pyroclasts have the largest range, from 50 to 1350 kg m-3, corresponding to vesicularities between 98% and 50%. While episode 45 pyroclasts have densities mainly between 600 and 1000 kg m-3, there is a small secondary peak between 400 and 500 kg m-3, corresponding closely to the modal densities and vesicularities of the other four episodes. 39

2.4.22.4 Results Microtextural analysis

2.4.2.1 Qualitative results

The pyroclasts from all five eruption episodes have highly variable vesicle microtextures that can be classified into three distinct categories. The texture within each studied pyroclast is indicated in Table 2.2. The ‘bubbly texture’ category is dominated by small, spherical vesicles (100-200 µm in diameter) that are separated by thick glass walls (commonly 50 µm thick, but up to 100 µm) (Fig. 2.5A). The round shape suggests that there has been minimal interaction between neighbouring bubbles. Clasts with this texture are typically the least vesicular and most dense. The ‘foamy texture’ category is dominated by large, closely packed polygonal vesicles, typically ≥500 µm, with bubble walls approximately 5 µm thick (Fig. 2.5B). The texture of these pyroclasts closely resembles reticulite texture as defined by Mangan and Cashman (1996), whereby the vesicles are polyhedral and are separated by micrometer-scale glass septa. Clasts with this texture have very high vesicularities, but very few fall within the vesicularity range for reticulite, defined as >98% by Mangan and Cashman (1996), and therefore we use the name foamy for this textural end-member. Although these two end-member textures differ greatly, they are not mutually exclusive, and there is a small population of pyroclasts that exhibit both textures (Fig. 2.5C). In such pyroclasts (only 3 of the total 19 in this study), the transition between bubbly and foamy domains is abrupt without transitional regions. Furthermore, there is no consistent spatial distribution of the two end-member textures in a pyroclast; for example, one end-member is not confined to the rim or core of a pyroclast. A third vesicle microtexture, which we call ‘irregular bubbly texture’, was found in one pyroclast from episode 32. Pyroclast 04B-38 has a texture similar to bubbly texture, in that the vesicles are separated by glass walls 100–250 µm thick. However, this pyroclast also has relatively large (1.00 and 1.24 mm) vesicles that are commonly polylobate or irregular in shape (Fig. 2.5D). Pyroclast 04B-38 is the only pyroclast within the studied sample set that contains observable microlites (Fig. 2.6).

2.4.2.2 Quantitative results

The vesicle number densities of the measured pyroclasts (Nv; number of vesicles per 40

Table2.4 Results 2.2 – Texture and vesicularity data for all analysed pyroclasts. m Pyroclast Episode Texture* Density Density-derived Nv N v VG/VL** (kg m-3) vesic. (%) 01-33 Ep. 45 B 960 67.3 5.2×105 1.6×106 2.06 01-08 F/B 840 71.4 6.1×105 2.2×106 2.50 01-06 F 620 79.0 1.5×105 7.3×105 3.76 01-01 F 460 84.6 3.9×104 2.5×105 5.49 02-04 Ep. 44 B 700 76.2 1.0×106 4.4×106 3.20 02-08 F 450 84.7 9.4×104 6.1×105 5.54 02-05 F/B 330 88.8 7.7×104 6.8×105 7.93 03-05 Ep. 40 B 770 74.1 8.9×105 3.5×106 2.86 03-07 B 520 82.3 6.8×105 3.8×106 4.65 03-17 F/B 360 87.8 7.9×105 6.5×106 7.20 03-08 F 260 91.2 3.1×104 3.5×105 10.36 04A-05 Ep. 37 B 940 68.3 1.3×106 4.2×106 2.15 04A-09 B 750 74.6 5.7×105 2.2×106 2.94 04A-01 F 330 88.7 3.7×104 3.2×105 7.85 04A-04 F 240 91.7 4.5×104 5.4×105 11.05 04B-20 Ep. 32 B 1,120 62.1 3.3×105 8.9×105 1.64 04B-38 IB 870 70.7 1.7×105 5.7×105 2.41 04B-16 F 350 88.3 5.6×104 4.8×105 7.55 04B-06 F 290 90.2 5.1×104 5.3×105 9.20 * F = Foamy, B = Bubbly, F/B = combination of both, IB = Irregular bubbly ** Vesicle to Melt Ratio unit volume of bulk rock) range by almost two orders of magnitude across all the episodes sampled, from 3.0×104 cm-3 in episode 40 (pyroclast 03-08) to 1.3×106 cm-3 in episode 37 (pyroclast

04A-05) (Table 2.2). Within each sample, Nv spans at least one order of magnitude; the largest range is in episode 37 pyroclasts, with a minimum of 3.7×104 cm-3 and a maximum of 1.3×106 cm-3. In order to account for the presence of the vesicle population itself, a melt correction is

m applied to the Nv values to obtain the vesicle volume density per unit melt, N v (Klug et al., 2002). After this correction, the range in vesicle number densities across all the episodes is from 2.5×105 in episode 45 (pyroclast 01-01) to 6.5×106 in episode 40 (pyroclast 03-17) (Table 2.2).

m The N v values for these five Hawaiian fountaining episodes at Puʻu ʻŌʻō are mostly less than those calculated for the 1959 eruption of Kīlauea Iki (Stovall et al., 2011; Stovall et al., 2012) and larger than those calculated for the 1969 eruption of Mauna Ulu (Parcheta et al., 2013) (Fig. 2.7). To better define the vesiculation histories during each eruptive episode, vesicle size 41

2.4 Results

Figure 2.5 – A) Binary image showing bubbly texture. B) Binary image showing foamy texture. C) Example of a pyroclast that exhibits both bubble and foamy texture. Note that that the transition between the two textural domains is sharp. D) Binary image showing irregular bubbly texture. distributions (VSDs) were constructed for each analysed pyroclast (Fig. 2.8). These graphs show the fraction of the total vesicle volume that is attributed to each vesicle size range, ex- pressed as equivalent diameter. In most cases the VSDs are unimodal and have a moderate range (0.04–1.24 mm) (Fig. 2.8). Within two of the five eruption episodes (episodes 45 and 37), there is a common trend of increasing unimodality and negative skew with respect to increasing pyroclast vesicularity for VSDs. This trend is defined by the migration of the median and modal vesicle size to larger sizes, coupled with a decrease in the number of vesicles ≤0.16 mm in size. The VSDs of episode 32 pyroclasts have a unique trend in that the pyroclasts have a consistent modal vesicle size between 0.4 and 0.6 mm. However, the proportion of vesicles ≤0.16 mm and ≥1.0 mm decreases and the proportion of vesicles between 0.5 and 0.6 mm increases, so that the modal vesicle size becomes more dominant and the size range narrows with increasing 42

2.4 Results

Figure 2.6 – Four separate 500x magnification images showing the microlite species, abundance and textures from pyroclast 04B-38. In all images cpx: clinopyroxene, P: plagioclase, mag: magnetite and ves: vesicle. See supplementary material for locations of images within the pyroclast.

Figure 2.7 – Plot showing vesicularity versus vesicle number density (Nmv) for three high Hawaiian fountaining eruptions at Kīlauea. 43 vesicularity2.5 Discussion (Fig. 2.8).

2.5 DISCUSSION

2.5.1 Determining syn-eruptive vesicle textures Within the incandescent fountain, the cooling rates of lapilli-sized clasts are predicted to be too slow to retain a signature of the magma properties at the time of fragmentation (Porritt et al., 2012). The residence time within the incandescent fountain results in a textural evolution from a relatively dense fluidal, clast to pumice, followed by quenching during transport in a fountain (Stovall et al., 2011, 2012). Stovall et al. (2011; 2012) showed that vesiculation

m stages can be identified by comparing the N v of each pyroclast to its vesicle-to-melt ratio (VG/

VL) as defined by Gardner et al. (1996). Each separate vesiculation process, such as nucleation, growth, coalescence and loss (and any feasible combination of these) produces a characteristic

m trend in both VG/VL versus N v plots and the associated VSDs (Shea et al., 2010; Stovall et al., 2011; Stovall et al., 2012). Following Stovall et al. (2011), we interpret pyroclasts that have

m relatively high N v and low VG/VL to resemble most closely the vesiculation conditions of the magma at fragmentation; we call these ‘primary’ pyroclasts. For all five episodes, these primary pyroclasts (hollow data points in Fig. 2.9A–C), which we interpret to approximate the syn-frag-

m mentation vesicle population, have bubbly textures, low VG/VL values and high N v (Fig. 2.9), and are likely to have been transported in the cooler outer margins of the jet (2. in Fig. 2.2), where rapid cooling of the pyroclast inhibited post-fragmentation vesiculation processes.

m 2.5.2 Relationship between vesicle number density (N v) and fountain height High volatile supersaturations associated with disequilibrium (rapid) magma ascent drives the rate of bubble nucleation, as demonstrated by experiments and numerical models (Parfitt, 2004; Toramaru, 1995; Vergniolle and Jaupart, 1990). Thus, bubble nucleation rates,

m and therefore the N v of the erupted tephra, can be used as a proxy for the volatile supersaturation and decompression rate of the magma. In the case of Hawaiian fountains, the maximum fountain height should, theoretically, be directly proportional to the degree of supersaturation

m and nucleation rate in the shallow ascending magma and the N v of the erupted tephra (Polacci et al., 2009; Stovall et al., 2012).

m In this study average N v for pyroclasts from each episode does not correlate with the 44

2.5 Discussion - (bubbles v m – Histograms showing vesi - showing Histograms – ) are listed in the top left. The in the top left. listed ) are -3 ), vesicularity (%) and density ), vesicularity -3 (kg cm vertical blue line through the 0.16 mm the through line blue vertical equivalent diameter is for comparison. Figure 2.8 Figure cle size distributions for representative pyroclasts from each high fountaining diameter Equivalent studied. episode is defined as the diameter for the with the same area as calculated of a circle Pyroclasts 2006). al., et (Adams vesicle from each episode are arranged top to bottom in order of increasing vesicularity (decreasing density). is giv number pyroclast the plot, each In N the and right top the in en cm 45 observed2.5 Discussion maximum fountain height (Fig. 2.10). In fact, pyroclasts from the highest fountain

m 5 -3 (episode 32, 391 m) have the lowest average N v (6.13×10 cm ). Figure 2.11 shows that the

m N v of the primary pyroclasts from Episodes 37, 40, 44 and 45 may be more closely related to

m fountain height than the average N v for each episode, and implies that the microtextural data from primary pyroclasts better represents the vesiculation processes that were occurring within the shallow conduit prior to fragmentation. However, the lowest VG/VL pyroclast from Episode

m 32 also has the lowest N v value despite having the highest peak fountain height of all the episodes. Of the primary pyroclasts found, pyroclast 02-04 from episode 44 and pyroclast 01-08 from episode 45 have quenched rinds and a vesicle size gradient from rind to interior (Fig.

2.11). These features suggest that, despite both pyroclasts having the lowest VG/VL values found for their respective episodes, even they underwent some post-fragmentation vesicle modification. Therefore, episodes 32, 44 and 45 contain no pyroclasts that truely exhibit pre-/syn- frag- mentation microtextures, althought the texture of the quenched rims preserves some indications of pre-/syn- eruptive conditions. Only pyroclasts 03-05 and 04A-05, from the episodes 40 and 37 samples respectively, appear to be unmodified by post-fragmentation vesiculation processes and to preserve the syn-fragmentation vesicle population. Both pyroclasts have bubbly textures in which larger and smaller vesicles define bands a few millimeters across Fig.( 2.11). Although it is not possible to quantitatively describe and compare the vesiculation histories and shallow conduit processes of these high fountaining episodes based solely on these two pyroclasts, each of their qualitative textures suggest that the ascending melt is thermally and (or) mechanically heterogeneous on a small scale during Hawaiian-style fountaining. Despite rigorous sampling techniques, pyroclasts with primary textures that describe the fluid dynamics of the magma near the time of fragmentation remain elusive due to the per- vasiveness of post-fragmentation vesiculation processes. As a result, it is not possible, based on

m N v, to quantitatively describe the fluid dynamics of the magma in the shallow conduit in such a way that allows us to make detailed comparisons between distinct fountaining episodes at a single vent or, more importantly, between fountaining episodes at different vents or different volcanoes. 46

2.5 Discussion

m Figure 2.9 – VG/VL versus N v plots for each eruption episode showing the post-fragmentation vesiculation process that is defined by the pyroclast microtextures (after Stovall et al., 2011). In all diagrams, open circles denote primary pyroclasts and the grey arrows highlight the associated trend. The images show the microtexture of the labelled pyroclasts that are representative of textures found for that episode. A) Episodes 37, 44 and 45, all of which show growth plus coalescence trends. B) Episode 40, which shows both a growth plus coalescence trend and a growth plus nucleation trend. C) Episode 32, which has a growth trend. D) Key showing the possible trends and their associated post fragmentation vesicle processes. L = Loss, C = Coalescence, G+C = Growth plus Coales- cence, G = Growth, G+N = Growth plus Nucleation, N = Nucleation. 47

2.5 Discussion m Figure 2.10 – Plot of average N v versus fountain height for each eruptive episode. Black m bars indicate the maximum and minimum N v. Numbers correspond to the episode.

2.5.3 Relationship between vesicle-to-melt ratio (VG/VL) and fountain height

2.5.3.1 Episodes 37, 44 and 45

Episodes 37, 44 and 45 (Fig. 2.9A) pyroclasts show trends associated with continued growth and coalescence of vesicles after fragmentation, most likely within the incandescent jet. For these episodes, continued bubble growth and coalescence transforms primary bubbly tex-

m tured pyroclasts (low VG/VL, high N v) into mature foamy textured pyroclasts (high VG/VL, low

m N v). Episodes 37 and 44 pyroclasts both have large ranges in VG/VL (ranges of approximately

9 and 5, respectively), whereas for episode 45 pyroclasts the range of VG/VL is relatively small (a change of approximately 3.5). The comparison of these episodes demonstrates that the magnitude of the change in VG/VL is proportional to fountain height and thus residence time of the pyroclast within the jet. Episodes 37 and 44 have maximum fountain heights of 352 and 308 m respectively, compared to 257 m for episode 45. We suggest that lapilli remained above the glass transition temperature within the relatively low jet of episode 45 for shorter times than during episodes 37 and 44. There was thus less time for post-fragmentation processes to modify the vesicle population in episode 45 before quenching, resulting in a large proportion of smaller vesicles, bubbly textures and a negatively skewed polymodal density distribution (Fig. 2.4).

Growth and coalescence of vesicles within the jet after fragmentation results in an in- crease in average vesicle size, decrease in number density, and a thinning of vesicle walls, ulti- 48 mately2.5 Discussion forming closely packed polyhedral vesicles (Mangan and Cashman, 1996). This vesicle evolution is reflected by the VSDs for pyroclasts from each episode, most notably when comparing episodes 45 and 37, which show the larger median and modal vesicle sizes, coupled with fewer vesicles ≤0.16 mm in diameter (Fig. 2.8). 2.5.3.2 Episode 32

The Episode 32 sample has a simple trend in Figure 2.9. In Figure 2.9 there is large

m change in VG/VL across the analysed pyroclasts with very little change in N v. The high VG/VL pyroclasts 04B-16 and 04B-06 have foamy textures that are consistent with continued growth of vesicles from bubbly textures (pyroclast 04B-20) in the jet after fragmentation (Fig. 2.9C).

The second lowest VG/VL pyroclast, 04B-38 has a unique texture defined by a wide range in vesicle sizes but dominated by vesicles that are between 0.62 and 2.47 mm (Fig. 2.8). Large vesicles are ragged or irregular in shape and separated by thick (100-400 µm) glass walls (Fig. 2.5D). Of the entire sample set described in this study, only pyroclast 04B-38 from episode 32 has microlites; its glass, on average, contains 5-10 areal percent microlites; some vesicle-poor regions contain up to 30 area percent microlites (Fig. 2.6). The qualitative textures of pyroclast

04B-38 suggest that it is, in fact, more evolved than both the foamy textures in the high VG/

VL pyroclasts (pyroclasts 04B-16 and 04-06) and the bubbly texture of the other low VG/VL pyroclast 04B-20. Large, irregular vesicles separated by thick glassy regions rich in microlites

m Figure 2.11 – A) Plot of primary pyroclast N v versus fountain height for each eruptive episode for which primary pyroclasts were identified. B) Qualitative microtexture of primary pyroclasts depicted inA). 49 are2.5 Discussionindicative of magmas that have undergone bubble coalescence, outgassing and microlite crystallisation prior to quenching (e.g., Polacci et al. 2006; Sable et al., 2009; Constantini et al., 2010). The source of this sub-population of dense, texturally evolved pyroclasts is unknown (Figs. 4, 5d). Their textures are commonly taken to reflect relatively long residence time before eruption (e.g., Cimarelli et al., 2010; Sable et al., 2006). The repose period before episode 32 was 38.4 days and shorter than the repose period before episode 40 (48.5 days), and yet episode 40 produced no pyroclasts with this texture. There are at least two possible sources for these texturally evolved pyroclasts: 1) magma that stalled in one of the subterranean reservoirs proposed by Garcia et al. (1992), or 2) lava that had previously ponded and drained back in the conduit (Greenland et al., 1988). Whatever their origin, it is unclear why these ragged-textured pyroclasts are present only in the deposits of episode 32 and not in any of the other four episodes.

2.5.3.3 Episode 40

Episode 40 has the most complex vesiculation history of all five episodes. The GV /VL

m vs. N v plot suggests that, after fragmentation, some pyroclasts underwent, bubble nucleation, growth, and coalescence (Fig. 2.9B). The growth plus coalescence trend is defined by pyroclasts 03-05, 03-07 and 03-08. Pyroclasts 03-05 and 03-07 are primary and have bubbly texture, while pyroclast 03-08 has foamy texture. These textures are similar to those that define the growth plus coalescence trends in episodes 37, 44 and 45.

2.5.4 Quantitative microtextural signatures as a proxy for fountain height

We define the difference between the maximumG V /VL value and minimum VG/VL value obtained for pyroclasts from each eruption episode as Δ(VG/VL). Δ(VG/VL) is used to quantify the degree of post-fragmentation vesicle growth that took place in order to modify pyroclasts from bubbly texture to more mature foamy texture. This can be used to infer the residence time of a pyroclast within the hot plume of the fountain above the glass transition temperature, which in turn, is dependent on the peak fountain height (Stovall et al., 2011; Stovall et al., 2012).

Figure 2.12 shows Δ(VG/VL) values for Episodes 32, 37, 40, 44 and 45 of the 1983-1986

Puʻu ʻŌʻō eruption as well as Δ(VG/VL) values for Episodes 1, 15 and 16 of the 1959 Kīlauea Iki eruption from (Stovall et al., 2011; Stovall et al., 2012), plotted against the peak fountain height for each respective episode. This plot shows that, for all eight of these selected Hawaiian 50

2.6 Conclusions Figure 2.12 – Plot showing Δ(VG/VL) vs peak fountain height for all fountaining episodes sampled in this study, as well as three episodes (1, 15 and 16) from the 1969 Kīlauea Iki eruption from Stovall et al. (2011) and Stovall et al. (2012). Linear regression is calculated on all data points.

fountains there is a good linear correlation between Δ(VG/VL) and peak fountain height. This correlation holds true whether for discrete fountaining episodes from a single vent or two separate eruptions from different vents. This correlation suggests that Δ(VG/VL) derived from the microtextural analysis of multiple clasts from the same narrow stratigraphic interval could be used to ascertain the approximate peak fountain height of unobserved Hawaiian fountaining eruptions.

2.6 CONCLUSIONS

The five high Hawaiian fountaining episodes from the Puʻu ʻŌʻō eruption presented in this study had similar vesicle populations before fragmentation, with bubbly textures and melt-

m 6 -3 6 -3 corrected vesicle volume density (N v) values between 2.2×10 cm and 6.5×10 cm . These

m N v values fill a gap between values previously recorded for high Hawaiian fountains at Mauna Ulu and Kīlauea Iki (Parcheta et al., 2013; Stovall et al., 2011; Stovall et al., 2012). However, the signature of decompression rate and magma ascent is heavily overprinted by post-fragmen-

m tation vesiculation processes. These processes result in a reduction in the N v of tephra from

m these episodes, except in the most primary pyroclasts defined by low GV /VL ratio and high N v. Growth plus coalescence is the dominant post-fragmentation vesiculation process that occurs within the jet of Hawaiian fountains. It has a profound effect on the vesicle microtextures to the 51

2.6 Conclusions m extent that the theoretical relationship N v, explosivity and fountain height in Hawaiian-style eruptions does not hold true. We propose that the value range between the maximum and minimum vesicle-to-melt ratio (Δ(VG/VL)) for the tephra of Hawaiian fountaining eruptions can be used as a proxy for peak fountain height. Chapter 3

Rootlesswater littoral cones in Hawaiʻi formed by sustained, confined mixing of lava and sea-

Authors

S.J. Holt, R.J. Carey, J. McPhie

3.1 INTRODUCTION Rootless cones are a common phenomenon in basalt-dominated volcanic provinces such as Hawaiʻi, Iceland, the Canary Islands, the Columbia River Basalt Province and Reunion Island (Clarke et al., 2005; Marie, 2002; Moore and Ault, 1965; Reynolds et al., 2015; Stearns and Macdonald, 1946; Thorarinsson, 1951, 1953). Rootless cones form when lava interacts with external water, resulting in explosive lava–water interactions and the formation of a rootless cone atop, or adjacent to, the feeder lava flow (Fisher and Schmincke, 1984; MacDonald, 1972; Moore and Ault, 1965; Thorarinsson, 1953; Thordarson and Höskuldsson, 2008; Thordarson et al., 1998). Rootless cones that are formed within the littoral zone are referred to here as rootless littoral cones. Hydrovolcanic explosions are driven by the mixing of lava and external water and are thought to be analogous to explosive Molten Fuel Coolant Interactions (MFCI) (Wohletz, 2002; Wohletz, 1983; Zimanowski et al., 2015; Zimanowski et al., 1997). Despite this common fundamental origin, rootless cone formation may involve different circumstances. For example, on Hawaiʻi it is common for lava to flow into the ocean and form single or clus- tered rootless cones at the coast, referred to as littoral cones (Brigham, 1909; Fisher, 1968; Mat- tox and Mangan, 1997; Orr, 2008). However, in Iceland, numerous rootless cones have formed in groups (also known singly in the lierature as pseudocraters) where tube-fed lava flowed over water-saturated sediments, such as river sediments or glacial outwash (Fagents and Thordarson, 53

2007;3.2 Historic Greeley rootless and Fagents, littoral 2001; cones Hamilton on Hawaiʻi et al., 2010; Stevenson et al., 2012; Thorarinsson, 1951, 1953; Thordarson and Höskuldsson, 2008). Based on the observed rootless littoral activity at Kamoamoa, Hawaiʻi between 1992 and 1994, Mattox and Mangan (1997) defined two lava–water mixing regimes for rootlees littoral cones; open and confined mixing. Open mixing occurs when lava flows into a body of water and lava–water interaction can freely occur (for example, when channelized ʻaʻā lava flows enter the ocean and when tube-fed lava enters the ocean following a bench collapse). Confined mixing occurs when tube-fed lava interacts with sea-water that rushes into the system via fractures in the tube wall. In Hawaiʻi, rootless littoral cones produced by open mixing have many features that are indicative of formation via explosive hydrovolcanic activity, i.e. high ash content, poor sorting, grainsize bimodality, equant blocky clasts and quench textures (Houghton and Hackett, 1984; Mastin, 2007; Mattox and Mangan, 1997; Walker, 1973; Wohletz, 1983). In comparison, confined mixing in Hawaiʻi has been observed to produce low-profile (8 m high) circular mounds of welded bombs and lapilli with minor amounts of ash (Mattox and Mangan, 1997); however, there are no reported examples of large-scale, well-preserved rootless littoral cone deposits attributed to confined mixing. Littoral cones at ʻAuʻau Point and Puʻu Kī (Fig. 3.1) on the south-western flank of Mau- na Loa, Hawaiʻi show none of the diagnostic deposit characteristics of explosive hydrovolcanic activity driven by open mixing of lava and water common to other historic rootless littoral cones on Hawaiʻi. The cone-forming units of these two rootless littoral cones are, in fact, more akin to those produced by ‘dry’ explosive activity found elsewhere on Hawaiʻi (Kauahikaua et al., 2003; Orr, 2013; Patrick et al., 2011). The aim of this study is to critically assess confined mixing as a model for genesis of the rootless littoral cones at Puʻu Kī and ʻAuʻau Point in order to build upon the deposit description and model of formation put forward by Jurado-Chichay et al. (1996a). We also consider the similarities and differences between Puʻu Kī South Cone and ʻAuʻau Point and Icelandic rootless littoral cones to highlight the implications these two cones have for criteria used to discriminate between ‘dry’ and explosive hydrovolcanic deposits.

3.2 HISTORIC ROOTLESS LITTORAL CONES ON HAWAIʻI On Hawaiʻi, on numerous occasions in recent history, the flow of lava into the ocean, the resulting hydrovolcanic explosive activity and the formation of rootless littoral cones have been directly observed and recorded (Brigham, 1909; Mattox and Mangan, 1997; Moore and Ault, 54

3.2 Historic rootless littoral cones on Hawaiʻi

Figure 3.1 – A: Digital elevation map of Puʻu Kī. The location of the outcrop shown in Figure 2 is indicated. On both maps, dotted lines represent the spatial extent of the cone structure and the inset images show the location of the cone on Hawaiʻi. Note: Blue corresponds to sea level. B: Digital elevation map of ʻAuʻau Point. Solid line represents escarpment encrusted with welded lapilli and bombs. The location of the outcrop shown in Figure 3 is indicated.

1965; Orr, 2008). Hydrovolcanic rootless littoral cones have formed at the ocean entries of both channelized ʻaʻa lava and tube-fed lava (Mattox and Mangan, 1997; Orr, 2008). However, lava– water interaction associated with channelized ʻaʻa lavas generates conditions that favour higher-intensity explosive hydrovolcanic activity compared with tube-fed lava (Jurado-Chichay et al., 1996a; MacDonald, 1972). The Hawaiian Volcano Observatory (HVO) has monitored lava entering the ocean at Kamoamoa (1992-1994), Waikupanaha (Apr. 2008-Mar. 2009) and on the recently active lava flow field of Kīlauea volcano (Mattox and Mangan, 1997; Orr, 2008). Mattox and Mangan (1997) defined four transient eruption mechanisms during the 1992-1994 explosive hydrovolcanic littoral activity of Kamoamoa and their associated products. (1) Tephra jets result from explosive interactions between lava and sea-water and are composed of black tephra and steam, reaching heights up to 100 m. Tephra jets have been the dominant eruption mechanism during observed hydrovolcanic rootless littoral activity (Mastin, 2007; Orr, 2008). Typical tephra jet products are fines dominated (80% <5mm) and consist of spatter bombs, lapilli and coarse ash. The ash fraction is dominated by Pele’s tears and Limu o Pele (Mattox and Mangan, 1997). (2) Lava bubble bursts are sporadic bursts of very large bubbles (up to 10 m across) of molten lava from skylights (also observed at Waikupanaha (Orr, 2008)) and their products include molten bombs and lapilli as well as minor amounts of limo-o-Pele and Pele’s tears (Mattox and Man- 55 gan,3.2 Historic 1997). (3) rootless Lithic blasts littoral are steamcones explosionson Hawaiʻi that eject fragments of previously emplaced, solid lava. Their products consist of lapilli- to block-sized lithic clasts. (4) Littoral lava fountains are the rarest eruption mechanism observed by Mattox and Mangan (1997) and comprise fountains of molten lava droplets, ash and steam up to 100 m high that have total durations of 10-45 minutes. Littoral lava fountains produce molten bombs and lapilli with minor amounts of ash (Mattox and Mangan, 1997). Mattox and Mangan (1997) attributed tephra jets and lithic blasts to an open mixing regime and bubble bursts and littoral lava fountains to a confined mixing regime. Although these four eruption mechanisms vary greatly in nature, they are all short-duration events. In the case of confined mixing, Mattox and Mangan (1997) attribute transitions from littoral lava fountaining and bubble bursting to variations in the rate of sea-water influx into the system, whereby, littoral lava fountains are the result of a steady, continuous influx of water and bubble bursts the result of unsteady infiltration of water into the lava tube. Rootless littoral cones that have formed during periods of observed explosive activity driven by open mixing share similar cone geometries and deposits. The cones have a characteristic half-cone geometry, as only the half-cone onshore is preserved (Fisher, 1968; Fisher and Schmincke, 1984; MacDonald, 1972; Walker, 1993). Where the rootless activity was fed by a channelized lava flow, the half-cone is bisected by the lava into two onshore quarter-cones, as any pyroclasts deposited on the lava flow surface are carried away (Fisher, 1968; Fisher and Schmincke, 1984; Jurado-Chichay et al., 1996a). The cones thin rapidly inland; they have linear thickness half-distances (T1/2, Houghton et al. 2004) of ~10 m and can range greatly in height from 10 m (Waikupanaha) to 35 m (Sand Hills). Half-cones are dominated by thinly planar-bedded to cross-bedded, bimodal deposits composed mainly of ash (70-90%) and minor lapilli and bombs. The ash component of rootless littoral cones formed by tube-fed lava, such as Waik- upanaha and Kalapana (Mattox and Mangan, 1997; Orr, 2008), consists of Pele’s hair, Pele’s tears, and Limu o Pele. Deposits associated with channelized ʻaʻa lava flows, such as Alika Bay (Moore and Ault, 1965), consist of ash that is composed predominantly of angular, blocky, glassy fragments lacking fluidal textures, and angular lapilli and blocks which typically have highly breadcrusted or quenched rims. In both cases (channelized and tube-fed lava flows), despite the difference in pyroclast shapes and textures, the deposits are poorly sorted and have bimodal grainsizes and high proportions of coarse ash. Such features are widely considered 56 to3.3 be Comparison diagnostic with of the Icelandic deposits rootlessfrom explosive cone groups hydrovolcanic activity (Mattox and Mangan, 1997; Valentine and Connor, 2015; Walker and Croasdale, 1971; Walker, 1992; Zimanowski et al., 2015).

3.3 COMPARISON WITH ICELANDIC ROOTLESS CONE GROUPS

Rootless cone groups in Iceland are formed when tube-fed lava flows over water-laden sediments, resulting in explosive lava–water inateractions (Thorarinsson, 1951, 1953). Rootless cone groups consist of 10s to 1000s of single cones that individually formed over periods of hours to days and that vary greatly in size, ranging from 5–450 m in diameter and 2–40 m high (Thordarson, 2000; Thordarson and Höskuldsson, 2008; Thordarson et al., 1998). Well studied examples of rootless cone groups include Leiðόlfsfell, Álftaver, Myvatn, Hnúta, Rauðhólar and Landbrot (Greeley and Fagents, 2001; Hamilton et al., 2017; Hamilton et al., 2010; Thordarson et al., 1998). Individual rootless cones within the cone groups can exhibit markedly different morphologies and deposit characteristics. Smaller diameter rootless cones tend to be steep sided, have small craters and have deposits that commonly consist of reversely graded, crudely bedded lapilli and coarse bombs that resemble spatter mounds or hornitos (Fagents et al., 2002). The larger diameter rootless cones have broad craters and resemble cinder cones or tuff rings (Fagents et al., 2002; Fagents and Thordarson, 2007). Rootless cones such as these are most common and have deposits that have distinct upper and lower sequences as defined by Fagents and Thordarson (2007) and Thordarson and Höskuldsson (2008). The lower sequence consists of well bedded, widely dispersed decimetre- to meter-scale beds of ash and lapilli that alternate with thin, crudely laminated beds of excavated substrate and black ash. The upper sequence is crudely bedded and consists of spatter bombs and agglutinate with a smaller proportion of substrate lithic clasts and ash and is capped by a 1–2 m thick unit of clastogenic lava. The upper sequence can also contain armoured bombs (also known as cored bombs) with a core usually consisting of lava crust (Fagents and Thordarson, 2007; Thorarinsson, 1951; Thordarson et al., 1998). Hamilton et al. (2010) further splits these deposits up into three main facies: 1) an ash rich distal deposit with sheet like geometry, likely deposited from a weak convective plume, 2) an ash and lapilli tephra platform consisting largely of fall deposits with rare dilute pyroclastic density current deposits, and 3) a vent-proximal lapilli and bombs that can be welded and ex- 57 hibit3.4 Rootless rheomorphic littoral flow. cones at Puʻu Kī and ʻAuʻau Point Four features that appear to be common to almost all Icelandic rootless cones are that their deposits are crudely- to well-bedded sometimes consisting of well-defined bed-pair successions, they contain abundant substrate lithic clasts (typically less abundant in the upper sequence than lower sequence), are poorly sorted at the bedding scale but are reversely graded over the entire deposit and are capped by a rheomorphic flow unit (Fagents and Thordarson, 2007; Fitch et al., 2017; Frey and Jarosewich, 1982; Hamilton et al., 2017; Hamilton et al., 2010; Thordarson et al., 1998). These deposit features indicate that the Icelandic rootless cones are formed from numerous, discrete explosions driven by dynamic physical mixing of saturated substrate and lava within a lava tube, and which decreased in eruption energy with time (Fa- gents et al., 2002; Fagents and Thordarson, 2007; Fitch et al., 2017; Frey and Jarosewich, 1982; Greeley and Fagents, 2001; Hamilton et al., 2017; Hamilton et al., 2010; Thordarson, 2000).

3.4 FIELD OBSERVATIONS OF ROOTLESS LITTORAL CONES AT PUʻU KĪ AND ʻAUʻAU POINT

3.4.1 Puʻu Kī South Cone

Puʻu Kī is located on the southwestern coast of Hawaiʻi, approximately 15 km northwest of Ka Lae (South Point). Puʻu Kī consists of two littoral cones, North Cone and South Cone (Fig. 3.1A), which both formed on the Pohūe Bay tube-fed lava (750-1500 years old) that was erupted from the Southwest Rift Zone of Mauna Loa (Jurado-Chichay and Rowland, 1995; Ju- rado-Chichay et al., 1996b; Wolfe, 1996). The cone geometry and deposit characteristics of the North Cone, such as high ash content and bimodal grainsize, suggest that it was formed through rootless activity (Jurado-Chichay et al., 1996a). This activity was likely driven by open mixing of lava and sea-water analogous to that described by Mattox and Mangan (1997). However, the South Cone is quasi-circular in plan view and has many deposit characteristics that are in striking contrast to those of the North Cone. We restrict our deposit description and discussion to Units A and B of the South Cone (Fig. 3.2) as the uppermost unit, Unit C, is a medial deposit from a period of explosive hydrovolcanic activity at the adjacent North Cone (Jurado-Chichay et al. 1996a). The nomenclature used here to describe the degree of welding follows that of Carey et al. (2008) and all componentry percentages are visual estimates. The oldest unit, Unit A, is also referred to as the Keliuli Cone (Jurado-Chichay et al., 58

1996a).3.4 Rootless Unit A littoral consists conesof thermally at Puʻu oxidised, Kī and ʻAuʻau slightly Point welded lapilli and coarse bombs (approximately 5-600 mm) and ≤5% ash component (Fig. 3.2), and shows weak decimetre-scale planar bedding, defined by grainsize and variations in the degree of welding. Unit B is well-bedded (bed thickness of 30-100 cm); the beds comprise alternating grey coherent basalt and thermally oxidised, slightly to moderately welded lapilli and bombs (10-250 mm) (Fig. 3.2). The coherent basalt beds are laterally continuous over distances up to 50 m and range in thickness from tens of centimetres to a metre. These coherent basalt intervals are interpreted as clastogenic lavas rather than primary lavas for three main reasons: (1) the coherent intervals grade vertically up and down into densely- to tack-welded deposits of bombs and fluidal lapilli Fig.( 3.3); there is also a vertical gradation, both up and down, in the degree of thermal oxidation that follows the degree of welding, from a non-oxidised grey coherent centre to strongly oxidised red, tack-welded pyroclasts; (2) the coherent intervals contain colour and vesicularity patterns that indicate cryptic outlines of remnant pyroclasts (Fig. 3.4), and (3) the coherent intervals mantle the underlying deposits, as would pyroclastic fall units (Fig. 3.2). Within Unit B, the pyroclastic beds between the clastogenic lavas are extensively thermally oxidised, tack- to moderately-welded lapilli and bomb layers containing 5% ash, are up to 4 m thick and are dominated by round cored bombs that are 8-10 cm in diameter.≪ These cored bombs have a core made of either a vesicular basalt clast or a dense, unoxidised basalt fragment, both with a rim of dense, commonly thermally oxidised basalt (Fig. 3.5).

Figure 3.2 – Puʻu Kī South Cone (location given in Figure 1A). Unit A: Keliuli Cone, Unit B: Inter- bedded clastogenic lavas (grey) and oxidised welded lapilli and bombs (red), Unit C: Unconsolidated ash and lapilli derived from Puʻu Kī North Cone. See text for detailed deposit descriptions. Sea cliff is approximately 20 m high. 59

3.4 Rootless littoral cones at Puʻu Kī and ʻAuʻau Point

Figure 3.3 – Clastogenic lava (grey) and thermally oxidised, welded lapilli and bombs from Puʻu Kī South Cone. Note that there is a gradation in welding away from the clastogenic lava. CL; Clastogenic lava, DW: Densely welded lapilli and bombs, W: Welded lapilli and bombs.

Figure 3.4 – A: A Remnant clast (indicated by arrow) within the coherent centre of a clastogenic lava lithofacies within Unit B at Puʻu Kī South Cone, also picture in Figure 3.3. B: Colour variations within the same lithofacies as A, which are defined by the thermal oxidation of remnant clasts.

3.4.2 ʻAuʻau Point

ʻAuʻau Point (Fig. 3.1B) is a rootless littoral cone that was fed by a pahoehoe branch of the Kolo lava flow. This lava was erupted from the Southwest Rift Zone of Mauna Loa between 200 and 750 years ago (Jurado-Chichay et al., 1996a; Wolfe, 1996). The cone is 30 m high and encloses an elongate 200-m-wide lava-filled crater Fig.( 3.1B). The cone appears to have origi- nally been circular, but the seaward side is cut by an escarpment that is ~10 m high. At the base of the escarpment is a lava bench. The escarpment is draped by coarse bombs. The cone at ʻAuʻau Point (Fig. 3.6) is largely built from Unit D, a 15-m-thick unit of alternating clastogenic lavas and thermally oxidised, slightly-welded lapilli and bombs, similar to 60

3.4 Rootless littoral cones at Puʻu Kī and ʻAuʻau Point

Figure 3.5 – A) An example of a cored bomb from Puʻu Kī South Cone. The core consists of an angular, vesicular basalt clast which is coated with a rim of dense, non-vesicular, thermally oxidised basalt. The core and rim appear to differ only in their vesicularity and level of thermal oxidation. B) A 4-5 m thick layer of tack- to moderately-welded, thermally oxidised lapilli and bombs from at Puʻu Kī South Cone. The visible, well-rounded bombs are predominantly cored bombs, similar to that pictured in A

Figure 3.6 – ʻAuʻau Point cone (location given in Figure 1B). Unit D: Interbedded clastogenic lavas (grey) and variably welded, oxidised lapilli and bombs (red), Unit E: non-welded ash-rich unit, Unit F: layer of coarse bombs. See text for detailed deposit descriptions. Sea cliff is approximately 30 m high. FZ = fault zone.

Unit B at Puʻu Kī South Cone. The welded lapilli and bomb layers in Unit D are 1 to 7 m thick and the clastogenic lavas have thicknesses up to 3 m. The T1/2 measured for this unit is approximately 30 m. Unit D is mantled by Unit E (Fig. 3.6), a non-welded, ash-rich unit approximately

15 m thick with a T1/2 of 40 m along the southern sea cliffs. Unit E ranges from planar strati- fied to cross-bedded and consists predominantly of coarse ash and lapilli (1-20 mm) with rare bombs (≥300 mm). The deposit characteristics of Unit E led Jurado-Chichay et al. (1996a) to suggest that it was produced by explosive hydrovolcanic activity, induced by seawater gaining 61 access3.5 Interpretation to molten lava of due deposit to a bench characteristics collapse. The uppermost unit at ʻAuʻau Point, Unit F (Fig. 3.6), consists solely of very large bombs (60-70 cm) on the inner wall and proximal outer walls of the cone. This unit was deposited simultaneously with lava effusion from a vent in the southwestern quadrant of the crater floor. The lava covered the floor of the crater and flowed over the escarpment down to sea level, forming the current surface of the lava bench. Approximately 70 m east and upslope from the eastern rim of ʻAuʻau Point cone there is a pit (Fig. 3.1B) which is approximately 80 m across and was described as a collapsed lava rise by Jurado-Chichay et al. (1996a).

3.5 INTERPRETATION OF DEPOSIT CHARACTERISTICS OF PUʻU KĪ SOUTH CONE AND ʻAUʻAU POINT The cone-forming units of the rootless littoral cones at ʻAuʻau Point and Puʻu Kī South Cone (Units A, B and D in Fig. 3.2 and 3.6 respectively) show many features traditionally attributed to ‘dry’ explosive magmatic activity, such as a high degree of sorting, the paucity of ash and quench textures, extensive thermal oxidation, a high degree of welding and fluidal pyroclast morphologies (Houghton and Gonnermann, 2008; Houghton and Hackett, 1984; Walker and Croasdale, 1971; Walker, 1993). Explosive rootless eruptions require an external energy source for fragmentation and ejection of pyroclasts other than the decompression of magmatic volatiles as they are sourced from lava flows that have undergone extensive outgassing (Sparks, 1978; Swanson and Fabbi, 1973) and they are not linked to the primary subterranean magmatic plumbing system of the volcano. Because the supply of magmatic volatiles is inadequate in this setting, it is not possible for these cones to be the result of ‘dry’ explosive magmatic activity, despite the lack of indicators suggesting explosive lava–water interaction. Assuming these cones are hydrovolcanic in origin, the only reasonable explanation for their apparently ‘dry’ deposit characteristics is for them to be formed from confined mixing as described by Mattox and Mangan (1997) but sustained for much longer periods and on a larger scale. The clastogenic lavas indicate periods of high discharge rate and high accumulation rate, such as littoral lava fountaining as defined by Mattox and Mangan (1997). A period of sustained littoral lava fountaining combined with high accumulation rates in the proximal region would account for the extensive welding and evidence for rheomorphic flow. The pyroclastic lithofacies of the cone-forming units consist of crudely-bedded, thermally oxidised tack- to moderately-welded lapilli and bombs. These lithofacies were likely formed during periods of 62 lower3.6 The discharge formation and of accumulation Puʻu Kī South rate Cone activity and ʻAuʻausuch as Point lava bubble bursts. Lower accumulation rates in the proximal zone produced thick units of variably-welded fluidal pyroclasts. During this time, numerous pyroclasts would not immediately escape the vent, and instead be recycled, generating the high proportion of cored bombs within the pyroclastic lithofacies. The circular geometry and characteristics of the cone-forming units at ʻAuʻau Point and Puʻu Kī South Cone set them apart from most other rootless littoral cones on Hawaiʻi (Jurado-Chichay et al., 1996a; Moore and Ault, 1965). Hawaiian rootless littoral cones are predominantly half- or quarter-cones that consist of thinly planar-bedded to cross-bedded, bimodal deposits composed mainly of ash (70-90%) and minor lapilli and bombs which are likely formed by lava water interaction in an open mixing regime (Mattox and Mangan, 1997). Prolonged confined mixing of lava and sea-water is the most likely mechanism that drove the formation of ʻAuʻau Point and Puʻu Kī South Cone, in a process similar to the formation of a small, steep-sided 7.5 m tall circular cone at Kamoamoa (Mattox and Mangan, 1997), but on a larger time-scale.

3.6 THE FORMATION OF PUʻU KĪ SOUTH CONE AND ʻAUʻAU POINT We propose that these two rootless cones formed as the result of hydrovolcanic rootless littoral activity driven by the confined mixing of lava and sea water. As the tube-fed lava flow advanced towards the ocean, the lava flowed over and interacted with the water-saturated substrate (Fig. 3.7A). The rootless eruption was then initiated when tensile fractures in the basal crust of the lava tube allowed the lava in the molten core of the tube to come into contact with sea-water (sourced from water-saturated substrate) and for physical mixing to occur (Fig. 3.7B) (Mattox and Mangan, 1997). The resulting rootless activity consisted of a period of sustained confined mixing of lava and sea-water, driving alternating periods of high discharge rate littoral lava fountaining (Fig. 3.7C) and periods of low discharge rate lava bubble bursting (Fig. 3.7D), as defined by Mattox and Mangan (1997). The alternating lithofacies within the cone-forming units of ʻAuʻau Point and Puʻu Kī South Cone reflect the fluctuating eruption style during the rootless activity. During periods of littoral lava fountaining (high energy high frequency explosions; Mattox and Mangan, 1997), high proximal accumulation rates resulted in the formation of clastogenic lavas (Fig. 3.7C). During periods of lower-intensity lava bubble bursts (low energy intermittent explosions; Mat- tox and Mangan, 1997) when water was depleted within the underlying substrate, lower prox- 63 imal3.7 Key accumulation Differences rates formed the thick, variably-welded lapilli and bomb lithofacies (Fig. 3.7D). The alternating lithofacies could reflect cycles of higher and lower eruption intensity and deposition rate over time. We propose that the alternating rootless eruption style and associated cone forming units at ʻAuʻau Point and Puʻu Kī South are the result of fluctuations in the lava–water mass ratio, driven by depletion and replenishment of available sea-water. Jurado-Chichay and Row- land (1995) showed that the lava flows that fed both ʻAuʻau Point and Puʻu Kī South Cone were likely to have had high lava fluxes (up to 890 3m s-1) and velocities (12 m s-1) within the lava tube, suggesting that any variations in water to lava mass ratio were not caused by a reduction in lava supply. The cycles in eruption intensity were therefore likely driven by fluctuations in the amount of water available to generate explosive hydrovolcanic activity. During periods of restricted, unsteady access of water, the lava–water mass ratio is unfavourable, resulting in bubble bursting. During periods of high water availability, the water to lava mass ratio was more ideal to promote extensive high intensity explosions driving littoral lava fountains. The eruption would then revert back to bubble bursting activity when the water within the substrate was locally depleted.

3.7 KEY DIFFERENCES BETWEEN PUʻU KĪ SOUTH CONE AND ʻAUʻAU POINT AND ICELANDIC ROOTLESS CONES ʻAuʻau Point and Puʻu Kī South Cone do not have any lithofacies with their cone-forming units that are diagnostic of explosive hydrovolcanic or MFCI activity, i.e. that are fine- grained, ash rich, lithic rich or poorly sorted. Lithofacies such as these at Icelandic rootless cones correspond to the lower sequence of pseudocraters defined by Fagents and Thordarson (2007), the distal sheet or tephra platform units defined by Hamilton et al. (2010) or the fine grained bed-pair and mud-rich layer successions observed at Rauðhólar by Hamilton et al. (2017). Therefore, the eruptions at ʻAuʻau Point or Puʻu Kī South Cone had much lower ex- plosivities than those in Iceland and at no point were there any high energy explosive pulses that produced dilute PDCs, like those described by Hamilton et al. (2010) and Hamilton et al. (2017), either at the onset of the eruption or at the onset of a discrete explosions during the eruption. In general the grain sizes of the cone-forming sequences of Icelandic rootless cones coarsen upwards over the entire deposit and represent a decrease in explosivity with time (Fa- 64

3.7 Key Differences

Figure 3.7 – Schematic of the formation of ʻAuʻau Point and Puʻu Kī South Cone. A) lava flows over saturated substrate. B: Tensile cracks form in the basal crust allowing water to percolate into the lava tube. 1. Surface crust 2. Water saturated substrate 3. Tensile fractures. C) Extensive mixing occurs between the sea-water and the molten core of the lava tube, initiating littoral lava fountaining and resulting in the deposition of clastogenic lavas. 4. Littoral lava fountaining activity 5. Clastogenic lavas 6. Steady influx of sea-water into the lava tube. D: The activity transitions into low energy lava bubble bursting activity. Heavy clast recycling within the vent results in lithofacies dominated by cored bombs. 7. Lava bubble bursting activity 8. Oxidised welded lapilli and bombs 9. Water level depletion, and reduction in water influx. gents and Thordarson, 2007; Hamilton et al., 2017; Hamilton et al., 2010) and they are typically capped by rheomorphic spatter (Fagents et al., 2002; Fagents and Thordarson, 2007; Frey and Jarosewich, 1982; Hamilton et al., 2010; Thordarson and Höskuldsson, 2008). In contrast, at ʻAuʻau Point and Puʻu Kī South Cone the alternating clastogenic lavas and slightly- to moderately-welded lapilli and bombs that comprise the cone-forming units suggest a cyclic eruption. The explosivity and deposition rates in the proximal zone increased and decreased multiple times throughout the eruption (fluctuating between littoral lava fountaining and bubble burst activity) rather than gradually waning from an initial high-intensity onset. The alternating clastogenic lava layers and welded pyroclastic layers at ʻAuʻau Point and Puʻu Kī South Cone could be considered analogous to the bed-pair layer successions at Rauðhólar (Hamilton et al., 2017) but are the product of lower intensity activity that underwent lower magnitude fluctuations in 65 explosivity3.8 Difficulties over inmuch Discriminating longer time scales.‘dry’ deposits from phreatomagmatic These critical differences in deposit characteristics reflect the fundamental difference in the driving mechanism of the eruptions at ʻAuʻau Point and Puʻu Kī South Cone versus Icelandic rootless cones. In the case of ʻAuʻau Point and Puʻu Kī South Cone, the eruption is driven by confined mixing of lava and sea-water and in the case of Icelandic rootless cones the eruption is driven by the physical mixing of lava and water-saturated substrate.

3.8 DIFFICULTIES IN DISCRIMINATING ‘DRY’ DEPOSITS FROM HYDROVOLCANIC DEPOSITS IN HAWAIʻI BASED ON DEPOSIT CHARACTERISTICS

Explosive rootless littoral activity in Hawaiʻi is driven by either open or confined mixing of lava and sea-water (Mattox and Mangan, 1997). These two mixing regimes can be identified based on field observations of the geometry of the rootless littoral cone produced and deposit characteristics. In general, open mixing produces rootless littoral cones with characteristic half- cone or quarter-cone geometries dominated by thinly planar-bedded to cross-bedded, bimodal deposits composed mainly of ash (70-90%) and minor lapilli and bombs. In contrast, confined mixing results in the formation of circular rootless littoral cones formed from the explosive interaction of the lava and sea-water within the lava tube, such as rootless littoral cones at Puʻu Kī South Cone and ʻAuʻau Point. Unlike open mixing, confined mixing can generate rootless cones that have deposit characteristics that can be both hydrovolcanic and ‘dry’ magmatic in appearance and texture. The identificationof rootless cones that have formed from explosive hydrovolcanic activity but have deposits that appear to be of ‘dry’ magmatic origin, such as Puʻu Kī South Cone, ʻAuʻau Point, raises questions about the diagnostic features that are used to distinguish between hydrovolcanic and ‘dry’ magmatic activity. Deposit characteristics such as a high degree of sorting, the paucity of ash and quench textures, extensive thermal oxidation, a high degree of welding and fluidal pyroclast morphologies suggest ‘dry’ magmatic activity despite their hydrovolcanic origins. In the modern context, interpretations such as these are not difficult because the context is preserved and provides important constraints. For example, cones formed at the terminus of lava flows are likely to be secondary, and cones formed in the littoral setting are likely to be hydrovolcanic in origin. However, in the ancient rock record, some rootless cones formed from confined mixing would be indistinguishable from cones around a primary vent formed from 66

‘dry’3.9 Conclusions magmatic activity due to the lack of geographical context. Furthermore, rootless cones formed from confined mixing are much more likely to be preserved in the ancient rock record than rootless littoral cones formed from open mixing, as their deposits are more consolidated and less prone to erosion. Therefore, rootless cones such as ʻAuʻau Point and Puʻu Kī South Cone present a specific challenge for volcanologists working in the ancient rock record, as the misidentification of rootless cones such as these as primary vents could result in large errors in the interpretation of ancient volcanoes magmatic plumbing systems.

3.9 CONCLUSIONS The characteristics of both ʻAuʻau Point and the Puʻu Kī South Cone differ greatly from observed examples of rootless littoral cones elsewhere on Hawaiʻi. Rootless cones on Hawaiʻi are predominantly half-cone or quarter-cone deposits dominated by thinly planar-bedded to cross-bedded, bimodal deposits composed mainly of ash (70-90%) and minor lapilli and bombs, formed from the open mixing of lava and sea-water as lava flows into the ocean. In contrast, these two rootless littoral cones consist of extensively thermally oxidised, non- to densely-welded fluidal lapilli and bombs, alternating with clastogenic lavas. These deposits are well sorted, ash-poor to ash-free. These two cones were formed by a sustained period of confined mixing between tube-fed lava and sea-water and are much larger than the deposits formed from observed examples of confined mixing on Hawaiʻi (Mattox and Mangan, 1997). The deposits of ʻAuʻau Point and Puʻu Kī South Cone highlight that not all rootless cones in Hawaiʻi are ob- viously hydrovolcanic. Prolonged confined mixing of lava and water can result in the formation of rootless cones that show very few of the diagnostic deposit characteristics of hydrovolcanic activity. Without geographical context, cones of this nature could easily be interpreted as ‘dry’ magmatic deposits at primary vents. Chapter 4

Passive outgassing processes of the lava lake at .

Kīlauea Volcano, Hawaiʻi

Authors

S.J. Holt, R.J. Carey, J. McPhie, J. Taddeucci, B.F. Houghton, T. Orr

4.1 INTRODUCTION

Long lived, stable lava lakes such as Erebus (Ilanko et al., 2015; Oppenheimer et al., 2009; Peters et al., 2014), Erta ‘Ale (Harris et al., 2005; Sawyer et al., 2008; Spampinato et al., 2008), Nyiragongo (Gerlach, 1980; Smets et al., 2016; Spampinato et al., 2013) and Kīlauea (Stovall et al., 2009; Orr and Rea, 2012; Orr et al., 2013; Patrick et al., 2016c; Gailler and Kaua- hikaua, 2017) provide opportunities to make direct observations of the upper reaches of a volcano’s magmatic system. The physical and chemical characteristics of lava lakes reflect processes that occur throughout the entire magmatic system, from the lava lake surface (Karlstrom and Manga, 2006; Oppenheimer and Yirgu, 2002; Orr et al., 2013), to within the lava lake (Nadeau et al., 2015; Patrick et al., 2016a; Patrick et al., 2016b; Smets et al., 2016), to shallow conduit (Harris et al., 2005; Oppenheimer et al., 2009; Vergniolle and Bouche, 2016) and deeper magma reservoir (Barnie et al., 2016; Patrick et al., 2015). The surfaces of lava lakes often consist of dark grey, viscoelastic, crustal plates that are separated by narrow hot incandescent sutures (Patrick et al., 2016a). The incandescent sutures often exhibit zig-zag rift patterns analogous to tectonic rift zones on the Earth’s surface (Duffield, 1972; Karlstrom and Manga, 2006). The incandescent plate boundaries are zones of crust formation whereby incandescent lava flows away from the boundary and cools to form dark grey crust (Patrick et al., 2016a). The flow direction and surface velocity of lava lakes, 68 manifested4.2 Lava lake as activiythe wholesale at Halemaʻumaʻu, movement Kīlaueaof crustal volcano plates across the lakes surface, has been intrinsically linked to the direction and speed of the convective flow of the hot magma within the lava lake below the viscoelastic crust (Harris et al., 2005). In addition to the conspicuous lateral movement of crustal plates, the level of the lava lake may fluctuate over short time-periods due to gas pistoning, which is thought to be driven by the build-up of volatiles in the upper levels of the lava lake and their subsequent release via large scale spattering at the lava-lake margin (Patrick et al., 2016b; Smets et al., 2016). Outgassing also takes place at the stable regions of the lava lake surface via small-scale transient events, such as bubble bursts and spot vents independent of gas-pistoning (Patrick et al., 2016b). The 2008-2018 eruption at the summit of Kīlauea Volcano began in March 2008 when an explosion opened a vent within Halemaʻumaʻu crater (Fig. 4.1), the first explosive activity since 1924 (Houghton et al., 2011; Wilson et al., 2008). From 2010 onwards there has been a stable lava lake within the Overlook Crater of Halemaʻumaʻu (Orr et al., 2013; Patrick et al., 2011). The combination of long-lived stability, its large size compared to other lava lakes and the extensive monitoring network and daily observations by the staff at the Hawaiian Volcano Observatory have allowed the lava lake at Kīlauea Volcano to be studied in detail using multiple techniques and datasets (Patrick et al., 2016a; Patrick et al., 2016b). This study utilises high-speed thermal infra-red videos and thermal image time-lapse sequences of the surface of Halemaʻumaʻu lava lake to document the small-scale, transient outgassing features such as bubble bursts and spot vents. By mapping the spatial distribution and temporal frequency of these features, we aim to determine whether these features are governed by shallow or deep-seated degassing processes, and the extent to which these small-scale features reflect the large-scale processes, such as magma ascent, degassing and convection, within the lava lake and/or shallow conduit of Kīlauea volcano’s summit.

4.2 LAVA LAKE ACTIVIY AT HALEMAʻUMAʻU, KĪLAUEA VOLCANO

4.2.1 Summary of the formation and evolution of Kīlauea’s lava lake

A detailed review of historical lava lake activity is given in Patrick et al. (2016a) and references there in. Kīlauea Volcano has had many summit lava lakes throughout its observed history, most of which have been located within Halemaʻumaʻu Crater (Patrick et al., 2016a). Prior to 2008, Kīlauea Volcano had two well-documented periods of near constant lava lake 69

4.2 Lava lake activiy at Halemaʻumaʻu, Kīlauea volcano

Figure 4.1 – A: Digital Elevation Map (DEM) of the Hawaiʻi. B: DEM of the summit region of Kīlauea Volcano, including the locations of the summit caldera, Halemaʻumaʻu, Overlook Vent and the East Rift Zone (ERZ). C: A close up of the lava lake within the Overlook Vent, which itself is located within Halemaʻumaʻu (image taken 9th December 2015). The dimensions of the lake are approximately 180 by 250 m. The dominant surface flow regime of the lava lake is from the NW Upwelling Zone (1) to the SW Sink (2) where spattering can be seen on the lake surface. activity, 1923 to1924 and 1967 to 1968 (Finch, 1947; Jaggar and Finch, 1924; Kinoshita et al., 1969; Patrick et al., 2016a). The most recent period of sustained activity at Kīlauea began on March 19th 2008 when a small explosion occurred simultaneously with the opening of a 35m-wide vent along the eastern rim of Halemaʻumaʻu (Wilson et al., 2008). This explosion was preceded by several months of increased seismic activity and SO2 emissions (Houghton et al., 2011; Wilson et al., 2008). Lava first became visible within the Overlook Crater in 2008 and continued sporadically 70 between4.2 Lava lake2008-2009 activiy as at ephemeral Halemaʻumaʻu, lava ponds Kīlauea (Orr volcano et al., 2013; Patrick et al., 2016b). By Feb- ruary 2010, a larger continuously active lava lake had appeared (Patrick et al., 2016b) and by 2011, had grown into a 150 metre-wide lava lake that extended into Halemaʻumaʻu (Orr et al., 2013). Over the next five years, the lava lake grew to its approximate current day dimensions of 180x250 m and the level of the lake surface has been on average between 30-60 m below the crater rim (Patrick et al., 2016a). The lava lake was stable within Halemaʻumaʻu until the 10th of May 2018 when it disappeared from view after nine days of continuous drop in the lava lake level due to the initiation of a new phase of eruptive activity in the lower East Rift Zone (Hawaiian Volcano Observatory, 2018; Liu et al., 2018; Neal et al., 2019).

4.2.2 Summary of gas-pistoning cycles and outgassing regimes of Kīlauea lava lake

Gas-pistoning activity was first described at Kīlauea by Swanson et al. (1979) at Mauna Ulu and then later at Puʻu ʻŌʻō on the East Rift Zone of Kīlauea by Orr and Rea (2012). Gas-pistoning is defined as the quiescent gradual rise of the lava lake surface, which is then abruptly terminated by simultaneous intense spattering and a sudden drop in the lava lake surface (Pat- rick et al., 2016b). Patrick et al. (2016b) showed that the lava lake at Halemaʻumaʻu was in a near continuous cycle of gas-pistoning during much of 2010 and 2013-2015, with an average of 2.4 events per day involving lava lake level changes of 11-22 m. Patrick et al. (2016b) concluded that the gas-pistoning cycles of the lava lake at Halemaʻumaʻu were controlled by shallow processes internal to the lava lake whereby a single cycle consisted of gases accumulating in a shallow foam layer (tens of metres deep) causing gradual lava level rise followed by sudden collapse and outgassing of the foam layer and associated lava level fall. Patrick et al. (2016b) also defined the two major outgassing regimes that are evident on the lava lake surface that are intrinsically linked with the stages of gas-pistoning: 1) Passive outgassing - the bursting of many small bubbles through the lava lake surface crust scattered around the lake, both at crustal plate boundaries and within plates. This outgassing regime accounts for approximately one third of SO2 emissions. 2) Spattering – where a train of larger metre scale bubbles bursts through the lava surface causing spattering, most commonly at the lake margins. These two outgassing regimes are used to define the overall behaviour of the lake where the ‘spattering regime’ (a combination of passive outgassing and spattering) which operates 71 when4.3 Datasets the lava lake level is “normal”, and periods of ‘non-spattering regime’ (exclusively passive outgassing) which operate when lava lake levels are elevated during gas-piston events (Patrick et al., 2016b).

4.2.3 Summary of surface flow regimes of Kīlauea lava lake

The surface of the lava lake at Halemaʻumaʻu generally flows from NW to SE (Fig. 4.2A), whereby the lava upwells along the northern margin of the lake and spreads towards the lake margins where there is downwelling. At the SE margin of the lake there is enhanced crustal plate downwelling at a point known as the ‘SE sink’ (Patrick et al., 2016a). The northern margin is generally quiescent, and characterised by transient metre-sized bubble bursts and no spattering, whereas downwelling at the SE sink is accompanied by spattering and vigorous surface plate consumption (Orr et al., 2014; Patrick et al., 2016a). Patrick et al. (2016a) split the surface flow of the lava lake at Halemaʻumaʻu into two general regimes. 1) Stable flow – whereby the lava lake was in a non-spattering regime or a spattering regime where spattering is occurring exclusively at the SE margin of lava with a steady-state flow direction and 2) Unstable flow – where spattering is elsewhere on the lake surface (other than the SE margin) resulting in a reduction in surface flow speed and a change in the surface flow vector away from steady-state. Patrick et al. (2016a) went on the define the driving mechanism behind the stable (steady-state) flow regime to be upwelling lava from deep within the lava lake and the driving mechanism behind unstable flow to be shallow spattering at the lava lake surface.

4.3 DATASETS

This study used two main datasets. High-speed thermal infrared videos were captured between the 2nd and the 6th of December 2013 from the old overlook on the rim of Halemaʻu- maʻu Crater directly above the lava lake within the Overlook Crater (Fig. 4.2B). The videos were captured using high-speed infra-red cameras at varying magnifications and frame rates. Relative temeperature profiles of the lava lake surface could be derived from the high-speed infra-red videos using FLIR software. All measured temperature values must be considered relative as the measurements were made through a shifting gaseous plume and thus do not ac- curactely represent the real temperature at the lava-lake surface. The second dataset comprises a near infrared time-lapse sequence of images of the lava 72

4.3 Datasets

Figure 4.2 – A: A far field view of the lava lake within the Overlook Vent showing the locations of the NW Upwelling Zone (NW UZ), the SW Sink as well as the general flow direction of the stable flow regime (arrow) as defined by (Patrick et al., 2016a). B: Infrared thermal images of stable flow regime (left) and unstable flow regime (right) from the HVO thermal camera. Arrow designates flow directions. C: Snapshots from high-speed infrared videos showing stable flow regime (left) and unstable flow regime (right). 73 lake4.4 Methods surface taken every second for 24 hours from 23:55 on the 1st of December 2013 to 23:55 on the 2nd of December 2013. These images were recorded by one of the permanent webcams deployed by the Hawaiian Volcano Observatory to continuously monitor Kīlauea Volcano (Pat- rick et al., 2014) (Fig. 4.2C). To complement these two detailed datasets, regular footage of the lava lake surface was captured from the 2nd to the 6th of December 2013 using a DLSR camera to record outgassing features that are otherwise invisible to the infra-red cameras.

4.4 METHODS Bubble burst distribution maps were generated by logging the location of bubble bursts by eye for both the high-speed thermal infrared videos and near infra-red time-lapse sequences. In the case of the time-lapse sequences, bubble bursts locations were logged on a frame by frame basis. Locations were logged for all intra-plate boundary bubble bursts and large, metre-scale boundary bubble bursts. Large boundary bubbles were only taken into account for two reasons: 1) In order to exclude the near constant bursting of centimetre sized bubbles at plate boundaries from the distribution and 2) at some times during the videos, the incandescence of the boundary itself could mask the thermal signatures of smaller bubble bursts, therefore only the large, clearly visible bubble bursts were logged. All lava lake surface observations in this study were recorded on video during periods of stable flow marked by the spattering flow regime (Fig. 4.2A), as defined by Patrick et al. (2016a). These conditions allowed the observation and analysis of small scale outgassing features that are present on the lake surface in regions where the viscoelastic crust is undisturbed by the spattering occurring at the SE sink during its steady state. Physical observations of passive outgassing features were interpreted from both high-speed thermal infrared videos and visible light videos to utilise the maximum resolution images.

4.5 RESULTS

4.5.1 Passive outgassing features

4.5.1.1 Discrete Bubble Bursts

Discrete bubble bursts are the most common passive outgassing feature seen on the lava lake surface at Halemaʻumaʻu. The sizes of the bubbles bursting at the lake surface vary 74 greatly4.5 Results from centimetres to several metres in diameter. Discrete bubble bursts occur at both incandescent boundaries between surface crust plates (boundary bubble bursts) or within plates (intra-plate bubble bursts). At all incandescent plate boundaries across the lake there is a near continuous bursting of centimetre sized bubbles at regular frequency, also observed by Patrick et al. (2016a). The single bubble bursts that constitute this activity cannot easily be discerned in either visible light or thermal images of the lake surface, however the process manifests as a constant flickering of incandescent light at plate boundaries in the high-speed infrared images. This small-scale bubble bursting at plate boundaries is often punctuated by the bursting of much larger (tens of centimetres to metre scale) bubble bursts which break through and dis- turb the boundary (Fig. 4.3). These large boundary bubble bursts are much less frequent than the high-frequency centimetre scale bubble bursting and can, in some cases, produce spatter that falls onto the lava lake surface. These large bubble bursts can be seen clearly in high-speed thermal infrared images as bright flashes of incandescence at plate boundaries. Intra-plate bubble bursts occur at a much lower frequency than the large boundary bubble bursts. Intra-plate bubble bursts occur where a large, typically ≥1 m, bubble impacts on the surface crust and bursts through (Fig. 4.4). The degree to which the bubble disturbs the crust can vary greatly. In some cases, the bubble is large enough and has enough energy to break entirely through, destroying the viscoelastic crust in a small region (Fig. 4.5). In other cases, a bubble causes the lake surface to rise locally and fracture the crust, but does not burst through, resulting in outgassing and ‘deflation’ of the bubble below the viscoelastic crust. Due to this variation in the extent of crust disturbance, intra-plate bubble bursts can appear much smaller on the thermal infrared images than they appear in visible light. Only the incandescent fractures in the crust are visible, but not the associated inflation(Fig. 4.5).

4.5.1.2 Surface Doming

We define surface doming events as instances where the surface crust of the lava lake inflates and begins to form a dome geometry but is not disrupted or fractured Fig.( 4.6). Sur- face doming events tend to be approximately 1-2 m in diameter. After inflation, the dome then subsides, sometimes resulting in short-lived radiating ripples on the crust surface. In instances where multiple surface domes form in quick succession, a fracture may appear in the crust resulting in the transition to successive discrete intra-plate bubble bursts. Due to there being no 75

4.5 Results

Figure 4.3 – Two separate time series, A-E and F-J showing the evolution of two examples boundary bubble bursts. In all cases the bubble bursts shown are metre scale and occur over 5-6 seconds. In each time series, the dashed boxes indicate static reference frames for each bubble burst. All bubble burst sequences occur over a time-scale of 1–2 seconds real time. 76

4.5 Results

Figure 4.4 – A sequence of frames showing a 4-5 m intra-plate bubble bursting event. White dotted lined indicates the approximate extent of the bubble. All bubble burst sequences occur over a time-scale of 1–2 seconds real time. incandescence during surface doming events, they are only readily discernible in visible light footage.

4.5.1.3 Spot Vents

Spot vents are the rarest types of passive outgassing observed on the lava lake surface. 77

4.5 Results

Figure 4.5 – Frames 1 to 6 show six separate intraplate bubble bursts (~1 m in size) in a fixed frame of reference on the lava lake surface. The frames span a time period of approximately three minutes. In each case the dotted line shows the approximate radius of the gas slug that impacts on the lake surface (inferred from previous frames). In each instance, the surface expression of the bubble burst and amount of incandescent lava revealed varies greatly despite being produced by gas slugs of similar size.

Spot vents are longer lasting features that can occur following an intra-plate bubble burst. They are characterised by ongoing puffs of magmatic volatiles out of the lake surface through a fracture or hole in the viscoelastic crust caused by an initial intra-plate bubble burst. Puffing associated with spot vents can continue for multiple seconds after the fracture is formed and the spot vent will migrate away from its initial point with the local plate motion, sometimes traversing a quarter of the lava lake’s length.

4.5.2 Distribution of discrete bubble bursts (boundary and intra-plate) The locations of discrete bubble burst events at the lava lake surface were mapped over two separate time periods, taking into account only intra-plate bubble bursts and large boundary bubble bursts. Firstly, bubble burst maps were generated for four separate high-speed thermal 78

4.5 Results

Figure 4.6 – An example of a surface doming event. Frame 1 shows the lava lake before event. Frame 2 shows the moment at which the slug (denoted by white arrow) impacts the lava lake crust, but does not break through. Frame 3 is the same moment in time as Frame 2, but has the extent of the surface doming event highlighted by white dashed lines.

infrared videos that were captured at various times throughout the day on December 3rd 2013 (Fig. 4.7). These videos correspond to 30 seconds to a minute real time and capture the majority of the lava lake surface area. They were recorded from the same location on the crater rim, however videos 20131203_001 and 20131203_003 were recorded at slightly higher magnifications (Fig. 4.7). Three of the four bubble burst maps generated (Fig. 4.7 A, B and C) show that discrete bubble bursts are localised along a diffuse elliptical ring centred around the north-western end of the lava lake which exists regardless of the plate configuration on the lake surface and is defined by both intra-plate and boundary bubble bursts. In addition, there are three separate zones along the ring (of very limited surface area marked on Figure 4.7), that show a much higher 79

4.5 Results

Figure 4.7 – Bubble burst distributions for four separate highspeed videos taken of the lava lake on De- cember 3rd 2016. Zones 1 to 3 are areas common to videos A, B and C which have the highest density of bubble bursts observed. These zones are not defined in video D asa it was much shorter in length, not allowing for enough bubble bursts to be captured. In all cases, annotation UZ denotes the NW Upwelling Zone. Note: temperature ranges are relative temperature. Definitive temperature measurements were not possible due to the masking effects of the ever-changing gas plume. frequency of bubble burst events. Relative temperature trends for Zone 1 from all three videos (Fig. 4.8) show that boundary bubble bursts occur at roughly three to five times the frequency of intra-plate bubble bursts within these zones. However, throughout the entire time period, bubble burst events occurred at a steady rate. The same bubble burst mapping method was applied to a time-lapse generated from thermal images taken of the lava lake every second for a period of 24 hours from 23:55 on the 1st of December 2013 to 23:55 on the 2nd of December 2013 (Fig. 4.9). The resolution of these images is much lower than the high-speed thermal infrared videos, meaning that only the largest magnitude bubble burst events were imaged and mapped. However, the much greater time period represented allowed a far greater number of bubble burst events to be recorded. This time-lapse shows that over longer time periods bubble burst events are again localised on an 80

4.5 Results

Figure 4.8 – A & B show two separate relative temperature profiles for Zone 1 from two videos, A and B in Figure 6. These plots show The Maximum (red) Average (grey) and Minimum (blue) relative temperatures for Zone 1. Inset images 1 & 2 show typical temperature profiles. Periods of elevated maximum temperature occur as a plate boundary passes through the reference frame, punctuated by high-frequency peaks. the high frequency peaks indicate brief exposure of incandescent lava associated with numerous boundary bubble bursts (red triangles). Periods of low maximum temperature occur when the interior of a plate occupies the reference frame. Low frequency peaks (green circles) indicate isolated intra-plate bubble bursts. The assymetrical peak represents the sudden exposure of incadescent lava, followed by gradual cooling. The amplitude of the intra-plate bubble burst peak is proportional to the degree of crust disturbance caused by the bubble (Figure 4.5). 81

4.6 Discussion

Figure 4.9 – Bubble burst distribution for a 24hr time-lapse video of the lava lake surface. Annotation UZ denotes the NW Upwelling Zone elliptical ring (albeit much less diffuse than the short high-speed thermal infrared video bubble burst maps) centred over the north-western end of the lake. Over the 24-hour period the three zones highlighted in Figure 4.7 become less pronounced due to the increased occurrence of bubble bursts elsewhere on the ring.

4.6 DISCUSSION

4.6.1 Relationships among passive outgassing processes

Although the passive outgassing processes observed in this study have very different physical characteristics (Fig. 4.10), they all appear to be caused by the impact of ascending gas slugs on the lava lake surface. The frequency of large bubble bursts on the lake surface, both boundary and intra-plate, suggests gas slugs continually arrive from depth, and impinge on the base of the viscoelastic crust (Fig. 4.10A & B). Although there is a clear difference in frequency between boundary bubble bursts and intra-plate bubble bursts (Fig. 4.8), we interpret that to be governed by the mechanical properties of the crust, not by a difference in the number of gas 82

4.6 Discussion

Figure 4.10 – Steps A to E show to formation and possible surface expressions of passive outgassing features of the lava lake at Kīlauea. Gas slugs rise through the lava within the lake (A) and impact on the viscoelastic crust at the lake surface (B) causing surface doming. If the slug does not have enough energy to break the crust, the crust rebounds (C & D) resulting in a surface doming event. If the slug possesses enough energy to break through the crust, a bubble burst occurs (E). If the crust is undisturbed, but fractures (F) gas is emitted and deflation occurs, resulting in a spot vent as the bubble moves along with surface plate motion (G). slugs rising from depth. The greater strength of the crust in the middle of plates, compared to plate boundaries (Karlstrom and Manga, 2006) means only the most energetic gas slugs break through. Observations of the initiation of both surface doming and spot vents indicate that they are both the results of gas slugs impacting on the viscoelastic crust, much like bubble bursts. In the case of surface doming, the slugs impact on the base of the crust but do not have the energy required to cause a fracture, resulting in crustal rise followed by rebound (Fig. 4.10B, C & D). In the case of spot vents, outgassing is initiated by a gas slug impinging on the surface and fracturing the viscoelastic layer, but not bursting through. The gas escapes through the newly formed ‘vent’ in the lava lake surface (Fig. 4.10F & G). Outgassing can then continue through this vent as the plate it occurs within moves across the lake surface, resulting in the observed gas puffs. However, it is unclear whether emitted gas was sourced from remnants of the original 83 gas4.6 Discussionslug, or whether it came from continued bursting of small bubbles within the foam layer beneath the crust.

4.6.2 Driving mechanism behind passive outgassing – Surficial or deep processes?

The surface features of lava lakes have been proven to be closely linked to surficial (the surface crust and uppermost underlying lava), shallow (within the lava lake) and deep (the volcanic conduit) processes (Harris et al., 2005; Patrick et al., 2016a; Patrick et al., 2015; Patrick et al., 2016b; Vergniolle and Bouche, 2016). Harris et al. (2005) concluded that the surface motion of lava lakes is driven by the internal convection of the lava and, therefore, the upper surfaces of lava lakes are mechanically coupled to the upper regions of the underlying magma. Patrick et al. (2016a) showed that the surface plate motion at Halemaʻumaʻu was driven by upwelling magma sourced deep within the lava lake, but the surface plates may be mechanically coupled to only the uppermost layer of the lava lake and thus decoupled from any convective flow deeper within the lava lake. Regardless of whether the plate motion is coupled to deep convection currents or only to the uppermost lava within the lake, features observed at the lava lake surface that occur at fixed points over time can therefore be interpreted to be decoupled from the internal convection within the lava lake, and are likely to be the result of deeper processes (Patrick et al., 2016a; Tazieff, 1994). Large bubble burst events at the surface of Halemaʻumaʻu lava lake are fixed spatially over a period of hours to days (Fig. 4.7 & 4.9). They are localised along a well-defined elliptical ring that remains stable over long periods of time (hours to days). This stability in position indicates that the source of the gas slugs that cause these bubble bursting events is decoupled from the surface flow regime of the lava lake as their impacts on the surface crust are spatially fixed, even though the crust is in constant motion. Therefore, we interpret these large bubble burst events (which account for the majority of passive outgassing events on the lava lake surface) to be sourced from deeper within the lava lake, i.e. from below the upper layer of magma that is mechanically coupled to the surface crust. It is unclear whether the motion of the surficial plates of the lake is decoupled from the internal convective regime (Patrick et al., 2016a), or intrinsically linked to the lake-wide convection (Harris et al., 2005). Microtextures of tephra erupted from the Overlook Vent during rockfall events suggest that lava lake convection takes place over the uppermost 100-300 m of the upper conduit at Kīlauea (Carey et al., 2013). The gas slugs that cause the bubble bursts 84

4.6 Discussion

Figure 4.11 – A: A schematic cross-section of the Kīlauea lava lake, showing the foamy layer below the viscoelastic crust and the assumed flow with the lake (arrows), from Patrick et. al (2016a). B: Sche- matic showing a possible intra-lava lake source for the slugs that drive passive outgassing at the lava lake surface. In this case the peripheries of the upwelling plume interact with the foamy layer causing either localised coalescence of bubbles or a preferential pathway for bubbles rise in a ring geometry. C: A schematic showing an upper volcanic conduit source for gas slugs that drive passive outgassing at the lava lake surface. In this case bi-directional flow at the opening of the conduit at the base of the lava lake causes shear induced coalescence of bubbles at the interface between rising bubbly magma and dense, degassed magma (in a bi-directional flow regime such as core-annular flow), Meter scale slugs then buoyantly ascend trough the lake decoupled from the lake convection. In both A and B grey dotted lines represent the approximate extent of upwelling magma, arrows represent flow direction. described in this study are sourced from below the zone of mechanically coupled lava and surface crust but the depth is unknown. They may be sourced from a region within the lava lake, or from deeper within the volcanic conduit. 85

4.6.34.6 Discussion Implications for deeper lava lake processes

Despite the paucity of information on the internal processes of stable lava lakes, the observation of fixed, long-lived zones of passive outgassing at the surface of Halemaʻumaʻu lava lake helps place constraints on the greater lava lake system. The occurrence of these bubble bursts suggests that there is a persistent gas supply to the lake surface and that metre scale gas slugs are formed and localised into a ring geometry, either in the shallow regions of the lava lake, or in the upper region of the conduit and rise through the lake largely decoupled from the overall convective regime. However, in either case, the slugs (localised in a ring geometry) are decoupled from the motion of the surface crust and uppermost magma. In this study we consider two potential mechanisms for the formation and localisation of the gas slugs that drive passive outgassing at Halemaʻumaʻu lava lake into a ring geometry (Fig. 4.10). In both hypoth- eses, the magma is assumed to rise through the conduit as a homogenous, bubbly, low-Reynolds number flow (Manga et al., 1998) until a critical point at which bubble coalescence occurs and metre-scale gas slugs are formed.

Intra-lava lake source

In this case, the metre-scale gas slugs are formed and localised into a ring geometry within the upper parts of the lake. Assuming an upwelling-driven flow regime (Patrick et al., 2016a), the ascending bubbly magma plume would pass through the foam layer that accumu- lates in the upper region of the lava lake (Patrick et al., 2016b; Smets et al., 2016). Numerical models indicate turbulence and shearing occurs at the peripheries of the upwelling plumes (List, 1982), which may lead to localised destabilization of the surrounding foam causing localised coalescence of bubbles or a preferential pathway for bubbles to rise in a ring geometry (Fig 4.11B).

Upper Volcanic Conduit Source In this case, the gas slugs that drive passive degassing at the lake surface are formed and localised into a ring geometry in the upper regions of the volcanic conduit (as it opens into the base of the lake) and rise through the lake largely undisturbed by the convective flow i.e. decoupled upwelling-driven flow regime defined by Patrick et al. (2016a) (Fig 4.11C). We propose that in this case the gas slugs are formed due to shear-induced coalescence of magmatic gas 86 bubbles4.6 Discussion as a result of bi-directional flow within the shallow conduit. Numerical models suggest that in order for a lava lake to remain active and stable for long periods of time there must be an input from depth of hot, low viscosity, volatile-rich, bubbly magma, which is accompanied by a return flow of dense, degassed magma down the feeder conduit in order to preserve mass within the lake (Witham and Llewellin, 2006). Bi-directional flow within the feeder conduit is a common feature in numerous models of the convective flow regimes within lava lakes (Har- ris et al., 2005; Oppenheimer et al., 2009; Patrick et al., 2016a). Bi-directional flow could, in theory, produce and localise gas slugs in a ring geometry due to the shear induced coalescence of bubbles (Caricchi et al., 2011; Okumura et al., 2009; Okumura et al., 2006) at the interface between bubbly magma and dense, degassed magma flowing in opposing directions in a bi-directional flow regime, such as core-annular flow (Beckett et al., 2011). It is most likely that this process takes place in the upper region of the conduit, where it opens into the base of the lava lake, as shear-induced coalescence and outgassing is most effective for pre-existing bubbles coarsened by mechanisms such as expansion and growth, i.e. higher up in the ascent pathway (Namiki, 2012).

Although the results of this study provide valuable insight into the surface expression of the underlying flow regimes within the lake, it does not provide any quantitative constraints on the nature of the source of the gas slugs that drive passive outgassing at the lake surface. In order to better understand whether the gas slugs are formed within the feeder conduit or within the lake, visual observations of outgassing phenomena during non-steady state events are required. Whether or not the static ring of bubble bursting events (and therefore the ring of slugs at depth) can withstand perturbations of the lava lake from steady-state, places constraints on the depth of the source of the gas slugs. For example, in the immediate period after large rockfalls, the upper region of lava lake undergoes large-scale upheaval and overturn, greatly disturbing the accumulated shallow foamy layer (Orr et al., 2013). How fast the ring of bubble bursts returns to the surface would indicate whether or not a shallow foamy layer is integral in its formation. 87

4.7 Conclusions CONCLUSIONS

Passive outgassing at the surface of Halemaʻumaʻu lava lake involves discrete bubble bursts, surface doming and spot vents. Surface doming and spot vents are closely related and occur where gas slugs are unable to break through the viscoelastic crust. Time-lapse and video observations show that bubble burst events at the surface of Halemaʻumaʻu lava lake are localised in a ring geometry that is static with time (hours to days) and independent of the overall motion of surface crust. Although these small-scale outgassing features do not provide any quantitative evidence of the interior convective flow regime of the lava lake, their fixed position indicates that they are mechanically decoupled from the surface motion of the lava lake. Chapter 5

Summary and Future Research

The three case studies in this project cover a wide range of examples of low intensity basaltic volcanism. Each example occurs in a very different volcanic environment, i.e., an open vent at Puʻu ʻŌʻō, an open vent capped by a stable lava lake at Halemaʻumaʻu lava lake, and rootless eruptions in the littoral setting at Puʻu Kī and ʻAuʻau Point. Despite occurring in very different volcanic environments, each example shows that complex parameters in the final stages of the transport are critical in defining the style of the eruption that takes place. Parameters such as volatile exsolution and vesiculation processes, conduit geometry and the interaction with external water, all play a key role in governing the eruption style. These three studies all provide valuable constraints on two key phenomena (conduit processes and dynamics, and eruption unsteadiness) identified by Bonadonna et al. (2016) as research priorities to help constrain the dynamics of volcanic eruptions.

5.1 CONDUIT PROCESSES AND DYNAMICS

5.1.1 Volatile exsolution and vesiculation processes

Studies have shown that the eruptive behaviour of basaltic volcanoes is heavily dependent on the depth, timing, and rate of volatile exsolution, and the buoyant ascent of large vesicles within the shallow conduit (section 2.1 and references therein). One of the main chal- lenges in trying to constrain these processes is understanding whether the vesicle populations preserved within pyroclasts are primary textures which record the vesicle population at the time of fragmentation, or whether they have been modified by post-fragmentation vesiculation processes. Results in Chapter 2 show that not all quantitative microtextural signatures within pyroclasts can be used to describe the vesicle population within the conduit pre-fragmentation.

m The vesicle volume density (N v) and vesicle size distribution (VSD) of pyroclasts studied cannot be assumed to be primary textures due to the extensive effects of post-fragmentation 89

5.1 Conduit Processes and Dynamics vesiculation processes. However, the new quantitative parameter Δ(VG/VL) (see section 2.5.4) can be utilised as both a proxy for the peak fountain height of Hawaiian fountaining eruptions as well as used as a tool to evaluate which pyroclast microtextures, if any within a collected sample, can be interpreted as primary. In order to make quantitative measurements of the depth, timing and rate of volatile exsolution during low intensity basaltic eruptions based on the microtexture of their deposits, it is critical that interpretations are made on pyroclasts that satisfy two criteria: 1) their measured

VG/VL is the lowest for the sample set collected for that eruptive episode, and 2) their qualitative microtextures show no signs of post-fragmentation modification, i.e., no vesicle size gradient from rind to core, no distinct zones of varying vesicle populations and no ragged coalesced vesicles. Using commonly accepted sampling techniques for microtextural analysis (see section 2.3) this study only yielded two pyroclasts that could be considered to have textures that met these two criteria. This method may not adequately account for the effects of post-fragmentation vesiculation processes. It is also important that sufficient primary pyroclasts are analysed to allow for quantitative comparisons across eruptive episodes and locations. Within this study only pyroclasts 03-05 and 04A-05, from the episodes 40 and 37 samples respectively, appear to be minimally modified by post-fragmentation vesiculation processes and preserve textures approaching the syn-fragmentation vesicle population. Both pyroclasts have bubbly textures with zones defined by bands of larger and smaller vesicles Fig.( 2.11). Although it is not possible to quantitatively describe and compare the volatile exsolution and vesiculation histories and shallow conduit processes of these high fountaining episodes based solely on these two pyroclasts, their qualitative textures suggest that the ascending melt is thermally and mechanically heterogeneous on a small scale during Hawaiian-style fountaining. In the case of Hawaiian fountaining the gas bubbles that drive the eruption are mechanically coupled to the ascending magma (Parfitt and Wilson, 1995). Observations presented in Chapter 4 show that a large proportion of the outgassing that takes place at the surface of the Halemaʻumaʻu lava lake is driven by gas slugs that are decoupled from the convection inferred by other studies (Patrick et al., 2016a) and rise independently due to their relative buoyancy. The ascent of the gas slugs is modulated such that they arrive at the surface of the lava lake along an elliptical ring geometry. The ring geometry is brought about in either the upper conduit due to variable magma flow regimes or within the lake itself by preferential ascent pathways 90 extending5.1 Conduit through Processes to the and uppermost Dynamics shallow foam layer beneath the crust (section 4.6.3). These observations can provide a valuable framework for future work to better understand the internal convection within the lava lake, inferred as the control which allows it to remain active rather than cool and solidify.

5.1.2 Magma/lava interaction with external water

The duration and intensity of phreatomagmatic eruptions are greatly affected by the extent to which molten magma or lava and external water interact and the conditions under which the interaction occurs. Rootless littoral cone eruptions provide great insights into the effects of variations in the interactions between lava and external water. Rootless littoral eruptions are short duration, transient phenomena that can vary in style and intensity over short periods of time, governed primarily by the nature of the lava–water interaction that occurs (Mattox and Mangan, 1997). Our observations at Puʻu Kī South Cone and ʻAuʻau Point support the findings of Mat- tox and Mangan (1997) that the environment of the lava-water interaction plays a large role in the explosivity, duration and style of the resulting rootless littoral eruption. Rootless littoral eruptions in Hawaiʻi that occur due to confined mixing of lava and water appear to produce low intensity littoral lava fountains and bubble bursts whereas those driven by open mixing consist of higher intensity tephra jets and lithic blasts. Furthermore, within each mixing regime (open and closed mixing) rapid and reversible changes in eruption style, i.e., changes from tephra jets to lithic blasts in the case of open mixing and littoral lava fountains to bubble bursts in confined mixing, appear to be a direct result of changes in the lava to water ratio during the interaction. The nature of the interaction not only controls the dominant eruption mechanism but also the deposit characteristics. In the case of the Hawaiian rootless littoral cones, open mixing between lava and water results in half- or quarter-cone deposits that have high ash contents, poor sorting, grainsize bimodality and quench-fragmented equant blocky clasts. In contrast, confined mixing can produce circular cones with welded pyroclastic deposits and extensive thermal oxidation, beds that have a high degree of sorting, a paucity of ash, and fluidal clast morphologies. These deposit characteristics are associated with confined mixing and suggest that under some conditions, the interaction between lava and water can yield deposits that have no diagnostics characteristics of phreatomagmatic activity and instead have characteristics usu- 91 ally5.2 Eruption attributed Unsteadiness to dry magmatic activity at primary vents.

5.2 ERUPTION UNSTEADINESS

Basaltic volcanic eruptions are often unsteady, whereby the intensity of the eruption can fluctuate over varying timescales. Understanding the processes that modulate eruption intensity is integral to the understanding of the behaviour of volcanic eruptions and the assessment of their associated hazards. The deposit characteristics of both Puʻu Kī South Cone and ʻAuʻau Point outlined in chapter 3 indicate that the eruptions that formed both of these cones were in- herently unsteady having repeated cycles featuring higher intensity littoral lava fountaining and lower intensity bubble bursts. At both locations the cycles had long periodicity where changes in intensity were not abrupt but gradual (indicated by the gradation in the welding of pyroclasts) (section 3.4 & 3.5). The cause of these fluctuations is the availability of water during the eruption and the resulting MFCI explosions. When the water is depleted, the eruption consists of low intensity bubble bursts and once the water is replenished the intensity of eruption increases to littoral lava fountaining until the water is depleted again, allowing the cycle to start over.

5.3 COMPARISON TO GLOBAL EXAMPLES AND FUTURE RESEARCH

Hawaiian fountaining

Microtextural analysis of pyroclasts from Hawaiian style fountaining remains a powerful tool that can be used to compare the deposits of separate fountaining episodes from a single location or eruptions from different volcanoes. However, the results of our study highlight that, due to the masking effects of post-fragmentation vesiculation processes, extreme caution must be taken in the comparison of datasets such as these if the qualitative and quantitative analytical methodologies are not the same in each case. This includes, not just the statistical calculations and image processing involved in the microtextural analysis, but also the criteria that are used to classify pyroclast vesicle textures as primary (see chapter 2).

Figure 2.12 in section 2.5 shows that the calculated values of Δ(VG/VL) scales with fountain height when compared across major historical examples of Hawaiian fountaining within the East Rift Zone of Kīlauea and can potentially be used as a proxy for fountain height when studying unobserved eruptions. It is critical however to expand the pyroclast microtextural data set that is the foundation of this relationship. The addition of microtextural data from the last 5.3 92 remainingComparison historical to Globaleruption Examples in the east andrift zoneFuture of Kīlauea, Research Mauna Ulu in 1969 (Parcheta et

m al., 2013) will allow for the Δ(VG/VL) to be calculated for high-vesicularity, low N v end-member pyroclasts (see Figure 2.7, secion 2.4). The inclusion of data from the Mauna Ulu eruption will allow for the Δ(VG/VL) vs. fountain height relationship proposed in this study to be tested across the entire spectrum of Hawaiian fountaining observed in the East Rift Zone of Kīlauea. It is also important to expand this investigation to include datasets from Hawaiian style fountaining eruptions from volcanoes outside of Hawaiʻi. Short duration, high-intensity lava fountaining eruptions are common at Mount Etna Volcano, Italy (Ganci et al., 2012; Polacci et al., 2009). Examples of these eruptions have key differences in the pyroclast microtexture (high microlite abundances) and conduit flow regime (annular flow vs. bubbly flow) when compared to examples from Kīlauea (Polacci et al., 2006; Vergniolle and Gaudemer, 2015). These key differences would make it possible to investigate the effects of complex magma properties and conduit flow regimes on the textural relationships proposed in this study. With this expanded dataset it would be possible to develop an improved, robust methodology for vesicle microtexture analysis that would ensure that pyroclasts with primary vesicle textures can be targeted. A greater number of analysed primary pyroclasts would allow for direct quantitative comparisons between differing Hawaiian-style fountaining eruptions in order to better understand the volatile exsolution and vesiculation processes that occur in the shallow conduit during the final stages of magma ascent.

Persistent Lava Lakes

Before any detailed comparisons can be made between the passive outgassing processes at Halemaʻumaʻu lava lake and other examples of persistent lava lakes, e.g. Erebus (Ilanko et al., 2015; Oppenheimer et al., 2009; Peters et al., 2014), Erta Ale (Harris et al., 2005; Sawyer et al., 2008; Spampinato et al., 2008) and Nyiragongo (Gerlach, 1980; Smets et al., 2016; Spampi- nato et al., 2013), further work needs to be carried out in order to constrain the longevity of the ring geometry described in chapter 4. This can be achieved by the addition of two key datasets. Firstly, by carrying out the bubble burst distribution mapping on a time-lapse sequence captured during multiple periods of stable flow regime before and after 2013 to validate whether this is a long-lasting, stable feature or simply a transient phenomenon observed during the field campaign in 2013. Secondly, further constraints on the depth at which the ring geometry is 5.3 93 formedComparison could be madeto Global by examining Examples the and lava Futurelake surface Research in the time period immediately after a rockfall has caused an unstable flow regime to develop. Large rockfalls can cause wholesale upheaval of the shallow foam layer and surface crust as well as initiating periods of unstable flow (Orr et al., 2013; Patrick et al., 2016a; Patrick et al., 2016b). If the ring geometry persists during upheaval and unstable flow, its stability would suggest that the ring is sourced from the upper parts of the shallow conduit, below the zone that is perturbed by the rockfall. Conversely if the ring geometry is destroyed during lake upheaval and slowly recovers over a time period of days, the recovery time may indicate that it is related to the shallow foam layer in the upper region of the lava lake as the ring geometry only returns once the foam layer has recovered post-upheaval.

Rootless littoral cones

Our observations at Puʻu Kī South Cone and ʻAuʻau Point suggest that these two cones are formed when tube-fed pahoehoe lava flows over a water-saturated substrate at, or below sea-level, resulting in eruptions driven by confined mixing of lava and sea-water that gained access to the lava tube. The cone-forming units of both ʻAuʻau Point and Puʻu Kī South Cone do not show any diagnostic characteristics of phreatomagmatism and are more akin to those produced by ‘dry’ magmatic activity and therefore show little resemblance to other rootless littoral cone deposits in Hawaiʻi and elsewhere. Further research is needed in order to ascertain the nature of the explosive interaction between the lava and sea-water that drives the high-intensity littoral fountaining activity, without producing the diagnostic signatures of MFCIs. Such research could concentrate on other known examples of circular rootless littoral cones in Ha- waiʻi, such as Three Rim Cone (Jurado-Chichay et al., 1996a), to ascertain how common littoral lava fountaining is at littoral cones and to better understand the associated deposits.

Bibliography

Abràmoff, M. D., Magalhães, P. J., and Ram, S. J., 2004, Image processing with ImageJ: Biophotonics international, v. 11, no. 7, p. 36-42. Adams, N. K., Houghton, B. F., and Hildreth, W., 2006, Abrupt transitions during sustained explosive eruptions: examples from the 1912 eruption of Novarupta, Alaska: Bulletin of volcanology, v. 69, no. 2, p. 189-206. Almendros, J., Chouet, B., Dawson, P., and Bond, T., 2002, Identifying elements of the plumbing system beneath Kilauea Volcano, Hawaii, from the source locations of very‐long‐period signals: Geo- physical Journal International, v. 148, no. 2, p. 303-312. Barberi, F., Bertagnini, A., Landi, P., and Principe, C., 1992, A review on phreatic eruptions and their precursors: Journal of volcanology and geothermal research, v. 52, no. 4, p. 231-246. Bargar, K., and Jackson, E., 1974, Calculated volumes of individual shield volcanoes along the Hawai- ian-Emperor chain: J. Res. US Geol. Surv, v. 2, no. 5, p. 545-550. Barnie, T., Oppenheimer, C., and Pagli, C., 2016, Does the lava lake of Erta ‘Ale volcano respond to regional magmatic and tectonic events? An investigation using Earth Observation data: Geolog- ical Society, London, Special Publications, v. 420, p. SP420. 415. Bonadonna, C., Cioni, R., Costa, A., Druitt, T., Phillips, J., Pioli, L., Andronico, D., Harris, A., Scollo, S., and Bachmann, O., 2016, MeMoVolc report on classification and dynamics of volcanic explosive eruptions: Bulletin of volcanology, v. 78, no. 11, p. 84. Borgia, A., 1994, Dynamic basis of volcanic spreading: Journal of Geophysical Research: Solid Earth (1978–2012), v. 99, no. B9, p. 17791-17804. Brigham, W. T., 1909, The Volcanoes of Kilauea and Mauna Loa on the Island of Hawaii: Their Various- ly Recorded History to the Present Time, Bishop Museum Press, v. 4. Brooks, B. A., Foster, J. H., Bevis, M., Frazer, L. N., Wolfe, C. J., and Behn, M., 2006, Periodic slow earthquakes on the flank of Kīlauea volcano, Hawaiʻi: Earth and Planetary Science Letters, v. 246, no. 3-4, p. 207-216. Carey, R. J., Houghton, B. F., and Thordarson, T., 2008, Contrasting styles of welding observed in the proximal Askja 1875 eruption deposits I: Regional welding: Journal of Volcanology and Geo- thermal Research, v. 171, no. 1–2, p. 1-19. Carey, R. J., Manga, M., Degruyter, W., Swanson, D., Houghton, B., Orr, T., and Patrick, M., 2012, Ex- ternally triggered renewed bubble nucleation in basaltic magma: The October 12 2008 eruption at Halema ‘uma ‘u Overlook vent, Kīlauea, Hawai ‘i, USA. Caricchi, L., Pommier, A., Pistone, M., Castro, J., Burgisser, A., and Perugini, D., 2011, Strain-induced magma degassing: insights from simple-shear experiments on bubble bearing melts: Bulletin of Volcanology, v. 73, no. 9, p. 1245-1257. Cashman, K. V., and Mangan, M. T., 1994, Physical aspects of magmatic degassing; II, Constraints on vesiculation processes from textural studies of eruptive products: Reviews in Mineralogy and Geochemistry, v. 30, no. 1, p. 447-478. Cashman, K. V., and Sparks, R. S. J., 2013, How volcanoes work: A 25 year perspective: Geological Society of America Bulletin, v. 125, no. 5-6, p. 664-690. 95

Cimarelli,Bibliography C., Di Traglia, F., and Taddeucci, J., 2010, Basaltic scoria textures from a zoned conduit as precursors to violent Strombolian activity: Geology, v. 38, no. 5, p. 439-442. Clague, D. A., and Dalrymple, G. B., 1987, The Hawaiian-Emperor volcanic chain. part I. Geologic evolution: Volcanism in Hawaii, v. 1, p. 5-54. Clarke, H., Troll, V., Carracedo, J., Byrne, K., and Gould, R., 2005, Changing eruptive styles and textural features from phreatomagmatic to strombolian activity of basaltic littoral cones: Los Erales cinder cone, Tenerife, Canary Islands: Estudios geológicos, v. 61, no. 3-6, p. 121-134. Costantini, L., Houghton, B., and Bonadonna, C., 2010, Constraints on eruption dynamics of basaltic explosive activity derived from chemical and microtextural study: the example of the Fontana Lapilli Plinian eruption, Nicaragua: Journal of Volcanology and Geothermal Research, v. 189, no. 3, p. 207-224. Crosson, R., and Endo, E., 1982, Focal mechanisms and locations of earthquakes in the vicinity of the 1975 Kalapana earthquake aftershock zone 1970–1979: Implications for tectonics of the south flank of Kilauea Volcano, island of Hawaii: Tectonics, v. 1, no. 6, p. 495-542. Decker, R., Koyanagi, R., Dvorak, J., Lockwood, J., Okamura, A., Yamashita, K., and Tanigawa, W., 1983, Seismicity and surface deformation of Mauna Loa volcano, Hawaii: Eos, Transactions American Geophysical Union, v. 64, no. 37, p. 545-547. Delaney, P. T., Fiske, R. S., Miklius, A., Okamura, A. T., and Sako, M. K., 1990, Deep magma body beneath the summit and rift zones of Kilauea Volcano, Hawaii: Science, v. 247, no. 4948, p. 1311-1316. Delaney, P. T., Miklius, A., Árnadóttir, T., Okamura, A. T., and Sako, M. K., 1993, Motion of Kilauea Volcano during sustained eruption from the Puu Oo and Kupaianaha vents, 1983–1991: Journal of Geophysical Research: Solid Earth (1978–2012), v. 98, no. B10, p. 17801-17820. Denlinger, R. P., and Okubo, P., 1995, Structure of the mobile south flank of Kilauea Volcano, Hawaii: Journal of Geophysical Research, v. 100, no. B12, p. 24499-24424,24507. Di Muro, A., Métrich, N., Allard, P., Aiuppa, A., Burton, M., Galle, B., and Staudacher, T., 2016, Mag- ma degassing at Piton de la Fournaise volcano, in Bacherley, P., Lenat, J.-F., Di Muro, A., and Michon, L., eds., Active Volcanoes of the Southwest Indian Ocean, Springer, p. 203-222. Duffield, W. A., 1975, Structure and origin of the Koae fault system, Kilauea volcano, Hawaii, US Gov- ernment Printing Office. Dvorak, J. J., Okamura, A. T., English, T. T., Koyanagi, R. Y., Nakata, J. S., Sako, M. K., Tanigawa, W. T., and Yamashita, K. M., 1986, Mechanical response of the south flank of Kilauea Volcano, Hawaii, to intrusive events along the rift systems: Tectonophysics, v. 124, no. 3, p. 193-209. Edmonds, M., and Gerlach, T. M., 2007, Vapor segregation and loss in basaltic melts: Geology, v. 35, no. 8, p. 751-754. Fagents, S., Lanagan, P., and Greeley, R., 2002, Rootless cones on Mars: a consequence of lava-ground ice interaction: Geological Society, London, Special Publications, v. 202, no. 1, p. 295-317. Fagents, S. A., and Thordarson, T., 2007, Rootless volcanic cones in Iceland and on Mars: The Geology of Mars: Evidence from Earth-Based Analog, v. 1, p. 151. Fedotov, S., Chirkov, A., Gusev, N., Kovalev, G., and Slezin, Y. B., 1980, The large fissure eruption in the region of Plosky Tolbachik volcano in Kamchatka, 1975–1976: Bulletin Volcanologique, v. 43, no. 1, p. 47-60. Fisher, R. V., 1968, Puu Hou littoral cones, Hawaii: Geologische Rundschau, v. 57, no. 3, p. 837-864. Fisher, R. V., and Schmincke, H.-U., 1984, Pyroclastic rocks, Springer-Verlag Berlin. Fitch, E. P., Fagents, S. A., Thordarson, T., and Hamilton, C. W., 2017, Fragmentation mechanisms associated with explosive lava–water interactions in a lacustrine environment: Bulletin of Vol- canology, v. 79, no. 1, p. 12. 96

Frey,Bibliography H., and Jarosewich, M., 1982, Subkilometer Martian volcanoes: Properties and possible terrestrial analogs: Journal of Geophysical Research: Solid Earth (1978–2012), v. 87, no. B12, p. 9867- 9879. Gailler, L., and Kauahikaua, J., 2017, Monitoring the cooling of the 1959 Kīlauea Iki lava lake using surface magnetic measurements: Bulletin of Volcanology, v. 79, no. 6, p. 40. Ganci, G., Harris, A. J., Del Negro, C., Guéhenneux, Y., Cappello, A., Labazuy, P., Calvari, S., and Gouhier, M., 2012, A year of lava fountaining at Etna: volumes from SEVIRI: Geophysical Research Letters, v. 39, no. 6. Garcia, M., 2017, How and Why Hawaiian Volcanism Has Become Pivotal to OUr Understanding of Volcanoes From Their Source to Surface, in Carey, R. J., Cayol, V., Poland, M., and Weis, D., eds., Hawaiian Volcanoes: From Source to Surface, Volume 208, John Wiley and Sons, Inc. Garcia, M. O., Rhodes, J. M., Wolfe, E., Ulrich, G. E., and Ho, R., 1992, Petrology of lavas from episodes 2–47 of the Puu Oo eruption of Kilauea Volcano, Hawaii: evaluation of magmatic processes: Bulletin of volcanology, v. 55, no. 1-2, p. 1-16. Gardner, J. E., Thomas, R. M., Jaupart, C., and Tait, S., 1996, Fragmentation of magma during Plinian volcanic eruptions: Bulletin of Volcanology, v. 58, no. 2-3, p. 144-162. Gerlach, T., 1980, Chemical characteristics of the volcanic gases from Nyiragongo lava lake and the generation of CH4-rich fluid inclusions in alkaline rocks: Journal of volcanology and geothermal research, v. 8, no. 2-4, p. 177-189. Gonnermann, H. M., and Manga, M., 2007, The fluid mechanics inside a volcano: Annu. Rev. Fluid Mech., v. 39, p. 321-356. Gonnermann, H. M., and Manga, M., 2013, Dynamics of magma ascent in the volcanic conduit, in Fa- gents, S., Gregg, T. K., and Lopes, R. M., eds., Modeling volcanic processes: the physics and mathematics of volcanism Cambridge University Press. Got, J., Fréchet, J., and Klein, F. W., 1994, Deep fault plane geometry inferred from multiplet relative relocation beneath the south flank of Kilauea: Journal of Geophysical Research: Solid Earth (1978–2012), v. 99, no. B8, p. 15375-15386. Greeley, R., and Fagents, S. A., 2001, Icelandic pseudocraters as analogs to some volcanic cones on Mars: Journal of Geophysical Research: Planets (1991–2012), v. 106, no. E9, p. 20527-20546. Greenland, L., Okamura, A., and Stokes, J., 1988, Constraints on the mechanics of eruption, in Wolfe, E., ed., The Puu O eruption of Kilauea Volcano, Episodes 1-20, January 3, 1983, to June 8, 1984, U.S. Government Printing Office. Guffanti, M., Diefenbach, A. K., Ewert, J. W., Ramsey, D. W., Cervelli, P. F., and Schilling, S. P., 2010, Volcano-monitoring instrumentation in the United States, 2008: US Geological Survey, 2331- 1258. Hamilton, C. W., Fitch, E. P., Fagents, S. A., and Thordarson, T., 2017, Rootless tephra stratigraphy and emplacement processes: Bulletin of Volcanology, v. 79, no. 1, p. 11. Hamilton, C. W., Thordarson, T., and Fagents, S. A., 2010, Explosive lava–water interactions I: architec- ture and emplacement chronology of volcanic rootless cone groups in the 1783–1784 Laki lava flow, Iceland: Bulletin of volcanology, v. 72, no. 4, p. 449-467. Harris, A. J., Carniel, R., and Jones, J., 2005, Identification of variable convective regimes at Erta Ale Lava Lake: Journal of Volcanology and Geothermal Research, v. 142, no. 3, p. 207-223. Hawaiian Volcano Observatory 2018, Preliminary summary of Kīlauea Volcano’s 2018 lower East Rift Zone eruption and summit collapse, United States Geological Society, viewed 02/12/2018, https://volcanoes.usgs.gov/vsc/file_mngr/file-192/PrelimSum_LERZ-Summit_2018.pdf Heliker, C., Kauahikaua, J., Sherrod, D. R., Lisowski, M., and Cervelli, P. F., 2003a, The rise and fall of 97

BibliographyPu ‘u ‘O ‘o cone, 1983–2002, in Heliker, C., Swanson, D., and Takahashi, T. J., eds., The Pu ‘u O ‘o-Kupaianaha Eruption of Kilauea Volcano, Hawaii: The First, Volume 20, U.S. Geological Survey, p. 29-51. Heliker, C., Swanson, D. A., and Takahashi, T. J., 2003b, The Pu’u ‘Ō’ō-Kūpaianaha Eruption of Kīlauea Volcano, Hawaiʻi: The First 20 Years, US Dept. of the Interior, US Geological Survey, v. 1676. Holcomb, R. T., Nelson, B. K., Reiners, P. W., and Sawyer, N.-L., 2000, Overlapping volcanoes: the origin of Hilo Ridge, Hawaii: Geology, v. 28, no. 6, p. 547-550. Houghton, B., and Gonnermann, H., 2008, Basaltic explosive volcanism: Constraints from deposits and models: Chemie der Erde-Geochemistry, v. 68, no. 2, p. 117-140. Houghton, B., and Hackett, W., 1984, Strombolian and phreatomagmatic deposits of Ohakune Craters, Ruapehu, New Zealand: a complex interaction between external water and rising basaltic magma: Journal of Volcanology and Geothermal Research, v. 21, no. 3, p. 207-231. Houghton, B., Swanson, D., Carey, R., Rausch, J., and Sutton, A., 2011, Pigeonholing pyroclasts: In- sights from the 19 March 2008 explosive eruption of Kīlauea volcano: Geology, v. 39, no. 3, p. 263-266. Houghton, B., and Wilson, C., 1989, A vesicularity index for pyroclastic deposits: Bulletin of volcanology, v. 51, no. 6, p. 451-462. Houghton, B., Wilson, C., Del Carlo, P., Coltelli, M., Sable, J., and Carey, R., 2004a, The influence of conduit processes on changes in style of basaltic Plinian eruptions: Tarawera 1886 and Etna 122 BC: Journal of Volcanology and Geothermal Research, v. 137, no. 1, p. 1-14. Houghton, B., Wilson, C., and Smith, I., 1999, Shallow-seated controls on styles of explosive basaltic volcanism: a case study from New Zealand: Journal of Volcanology and Geothermal Research, v. 91, no. 1, p. 97-120. Houghton, B. F., Wilson, C., Fierstein, J., and Hildreth, W., 2004b, Complex proximal deposition during the Plinian eruptions of 1912 at Novarupta, Alaska: Bulletin of volcanology, v. 66, no. 2, p. 95- 133. Ilanko, T., Oppenheimer, C., Burgisser, A., and Kyle, P., 2015, Transient degassing events at the lava lake of Erebus volcano, Antarctica: Chemistry and mechanisms: GeoResJ, v. 7, p. 43-58. Jaggar, T. A., and Finch, R. H., 1924, The explosive eruption of Kilauea in Hawaii, 1924: American Journal of Science, no. 47, p. 353-374. Johnson, M. C., Anderson, A. T., and Rutherford, M. J., 1994, Pre-eruptive volatile contents of magmas: Reviews in Mineralogy and Geochemistry, v. 30, no. 1, p. 281-330. Jurado-Chichay, Z., and Rowland, S. K., 1995, Channel overflows of the Pōhue Bay flow, Mauna Loa, Hawai’i: examples of the contrast between surface and interior lava: Bulletin of volcanology, v. 57, no. 2, p. 117-126. Jurado-Chichay, Z., Rowland, S. K., and Walker, G. P., 1996a, The formation of circular littoral cones from tube-fed pāhoehoe: Mauna Loa, Hawai’i: Bulletin of volcanology, v. 57, no. 7, p. 471-482. Jurado-Chichay, Z., Urrutia-Fucugauchi, J., and Rowland, S. K., 1996b, A paleomagnetic study of the Pohue Bay flow and its associated coastal cones, Mauna Loa volcano, Hawaii: constraints on their origin and temporal relationships: Physics of the Earth and Planetary Interiors, v. 97, no. 1, p. 269-277. Kauahikaua, J., Sherrod, D. R., Cashman, K. V., Heliker, C., Hon, K., Mattox, T., and Johnson, J., 2003, Hawaiian lava-flow dynamics during the Puu Oo-Kupaianaha eruption: A tale of two decades: The Puu Oo-Kupaianaha Eruption of Kilauea Volcano, Hawaii, the First, v. 20, p. 63-87. Karlstrom, L., and Manga, M., 2006, Origins and implications of zigzag rift patterns on lava lakes: Jour- nal of volcanology and geothermal research, v. 154, no. 3, p. 317-324. Klein, F., Koyanagi, R., Nakata, J., and Tanigawa, W., 1987, The seismicity of Kilauea’s magma system: 98

BibliographyUS Geol. Surv. Prof. Pap, v. 1350, no. 2, p. 1019-1185. Klug, C., Cashman, K., and Bacon, C., 2002, Structure and physical characteristics of pumice from the climactic eruption of Mount Mazama (Crater Lake), Oregon: Bulletin of Volcanology, v. 64, no. 7, p. 486-501. Koyaguchi, T., and Woods, A. W., 1996, On the formation of eruption columns following explosive mixing of magma and surface‐water: Journal of Geophysical Research: Solid Earth, v. 101, no. B3, p. 5561-5574. Koyanagi, R. Y., Krivoy, H. L., and Okamura, A. T., 1966, The 1962 Kaoiki, Hawaii, earthquake and its aftershocks: Bulletin of the Seismological Society of America, v. 56, no. 6, p. 1317-1335. Langenheim, V. A., and Clague, D. A., 1987, The Hawaiian-Emperor volcanic chain, Part II; Stratigraph- ic framework of volcanic rocks of the Hawaiian Islands: Volcanism in Hawaii, v. 1, p. 55-84. Lautze, N. C., and Houghton, B. F., 2005, Physical mingling of magma and complex eruption dynamics in the shallow conduit at Stromboli volcano, Italy: Geology, v. 33, no. 5, p. 425-428. Lautze, N. C., and Houghton, B. F., 2007, Linking variable explosion style and magma textures during 2002 at Stromboli volcano, Italy: Bulletin of Volcanology, v. 69, no. 4, p. 445-460. Lipman, P., 1980a, Rates of volcanic activity along the southwest rift zone of Mauna Loa volcano, Ha- waii: Bulletin of Volcanology, v. 43, no. 4, p. 703-725. Lipman, P. W., 1980b, The southwest rift zone of Mauna Loa: Implications for structural evolution of Hawaiian volcanoes: American Journal of Science, v. 280, p. 752-776. Lipman, P. W., Lockwood, J., Okamura, R., Swanson, D. t., and Yamashita, K., 1985, Ground deformation associated with the 1975 magnitude-7.2 earthquake and resulting changes in activity of Kilauea volcano, Hawaii. Lipman, P. W., Rhodes, J., and Dalrymple, G. B., 1990, The Ninole Basalt—implications for the structural evolution of Mauna Loa volcano, Hawaii: Bulletin of Volcanology, v. 53, no. 1, p. 1-19. Lipman, P. W., and Swenson, A., 1984, Generalized geologic map of the southwest rift zone of Mauna Loa Volcano, Hawaii. Liu, C., Lay, T., and Xiong, X., 2018, Rupture in the 4 May 2018 MW 6.9 Earthquake Seaward of the Kilauea East Rift Zone Fissure Eruption in Hawaii: Geophysical Research Letters, v. 45, no. 18, p. 9508-9515. Llewellin, E., and Manga, M., 2005, Bubble suspension rheology and implications for conduit flow: Journal of Volcanology and Geothermal Research, v. 143, no. 1, p. 205-217. Lockwood, J. P., and Lipman, P. W., 1987, Holocene eruptive history of Mauna Loa volcano: US Geol. Surv. Prof. Pap, v. 1350, no. 1, p. 509-535. Lundgren, P., Poland, M., Miklius, A., Orr, T., Yun, S. H., Fielding, E., Liu, Z., Tanaka, A., Szeliga, W., and Hensley, S., 2013, Evolution of dike opening during the March 2011 Kamoamoa fissure eruption, Kīlauea Volcano, Hawaii: Journal of Geophysical Research: Solid Earth. MacDonald, G. A., 1972, Volcanoes, University of California, Prentice Hall, 510 p.: Manga, M., Castro, J., Cashman, K. V., and Loewenberg, M., 1998, Rheology of bubble-bearing magmas: Journal of Volcanology and Geothermal Research, v. 87, no. 1, p. 15-28. Mangan, M. T., and Cashman, K. V., 1996, The structure of basaltic scoria and reticulite and inferences for vesiculation, foam formation, and fragmentation in lava fountains: Journal of Volcanology and Geothermal Research, v. 73, no. 1, p. 1-18. Marie, G., 2002, Les deltas de lave du littoral hawaїen: Revue de L’Association Volcanologique Eu- rpéenne, v. 95, p. 13-19. Mastin, L. G., 2007, Generation of fine hydromagmatic ash by growth and disintegration of glassy rinds: Journal of Geophysical Research: Solid Earth (1978–2012), v. 112, no. B2. 99

Mattox,Bibliography T. N., and Mangan, M. T., 1997, Littoral hydrovolcanic explosions: a case study of lava-seawater interaction at Kilauea volcano: Journal of Volcanology and Geothermal Research, v. 75, no. 1-2, p. 1-17. Miklius, A., Lisowski, M., Delaney, P. T., Denlinger, R. P., Dvorak, J. J., Okamura, A. T., and Sakol, M. K., 1995, Recent inflation and flank movement of Mauna Loa volcano: Mauna Loa Revealed: Structure, Composition, History, and Hazards, p. 199-205. Montagna, C. P., and Gonnermann, H. M., 2013, Magma flow between summit and Pu’u’Ō’ō at Kīlauea Volcano, Hawai’i: Geochemistry, Geophysics, Geosystems. Moore, J. G., and Ault, W. U., 1965, Historic littoral cones in Hawaii. Moore, J. G., and Clague, D. A., 1992, Volcano growth and evolution of the island of Hawaii: Geologi- cal Society of America Bulletin, v. 104, no. 11, p. 1471-1484. Moore, R. B., Helz, R. T., Dzurisin, D., Eaton, G. P., Koyanagi, R. Y., Lipman, P. W., Lockwood, J. P., and Puniwai, G. S., 1980, The 1977 eruption of Kilauea volcano, Hawaii: Journal of Volcanol- ogy and Geothermal Research, v. 7, no. 3, p. 189-210. Morgan, J. K., Park, J., and Zelt, C. A., 2010, Rift zone abandonment and reconfiguration in Hawaii: Mauna Loa’s Ninole rift zone: Geology, v. 38, no. 5, p. 471-474. Nadeau, P. A., Werner, C. A., Waite, G. P., Carn, S. A., Brewer, I. D., Elias, T., Sutton, A. J., and Kern, C., 2015, Using SO2 camera imagery to examine degassing and gas accumulation at Kilauea volcano, May 2010: J Volcanol Geotherm Res, v. 300. Namiki, A., 2012, An empirical scaling of shear-induced outgassing during magma ascent: Intermittent magma ascent causes effective outgassing: Earth and Planetary Science Letters, v. 353, p. 72-81. Neal, C., Brantley, S., Antolik, L., Babb, J., Burgess, M., Calles, K., Cappos, M., Chang, J., Conway, S., and Desmither, L., 2019, The 2018 rift eruption and summit collapse of Kīlauea Volcano: Science, v. 363, no. 6425, p. 367-374. Neal, C. A., and Lockwood, J. P., 2003, Geologic Map of the Summit Region of Kilauea Volcano, Ha- waii: USGS, scale 1:24000. Okumura, S., Nakamura, M., Takeuchi, S., Tsuchiyama, A., Nakano, T., and Uesugi, K., 2009, Mag- ma deformation may induce non-explosive volcanism via degassing through bubble networks: Earth and Planetary Science Letters, v. 281, no. 3, p. 267-274. Okumura, S., Nakamura, M., and Tsuchiyama, A., 2006, Shear-induced bubble coalescence in rhyolitic melts with low vesicularity: Geophysical Research Letters, v. 33, no. 20. Okubo, P. G., Benz, H. M., and Chouet, B. A., 1997, Imaging the crustal magma sources beneath Mauna Loa and Kilauea volcanoes, Hawaii: Geology, v. 25, no. 10, p. 867-870. Oppenheimer, C., Lomakina, A. S., Kyle, P. R., Kingsbury, N. G., and Boichu, M., 2009, Pulsatory magma supply to a phonolite lava lake: Earth and Planetary Science Letters, v. 284, no. 3, p. 392-398. Oppenheimer, C., and Yirgu, G., 2002, Thermal imaging of an active lava lake: Erta'Ale volcano, Ethi- opia: International Journal of Remote Sensing, v. 23, no. 22, p. 4777-4782. Orr, T., 2008, Hawaiian Volcano Observatory Quarterly Reports - 2008-2009 Waikupanaha notes. Orr, T., 2008, Low intensity hawaiian fountaining as exemplified by the March 2011, Kamoamoa eruption at Kilauea Volcano, Hawaiʻi, in Proceedings AGU Fall Meeting, San Francisco, 2013. Orr, T. R., Poland, M. P., Patrick, M. R., Thelen, W. A., Sutton, A. J., Elias, T., Thornber, C. R., Parcheta, C., and Wooten, K. M., 2015, Kılauea’s 5–9 March 2011 Kamoamoa fissure eruption and its relation to 30+ years of activity from Pu’u’O̅ ’o: Hawaiian Volcanoes: From Source to Surface, v. 208, p. 393. Orr, T. R., Thelen, W. A., Patrick, M. R., Swanson, D. A., and Wilson, D. C., 2013, Explosive eruptions 100

Bibliographytriggered by rockfalls at Kīlauea volcano, Hawai ‘i: Geology, v. 41, no. 2, p. 207-210. Papale, P., 1998, Volcanic conduit dynamics, in Freundt, A., and Rosi, M., eds., From magma to tephra: modelling physical processes of explosive volcanic eruptions, Elsevier Science, p. 55-91. Papale, P., 2005, Determination of total H2O and CO2 budgets in evolving magmas from melt inclusion data: Journal of Geophysical Research: Solid Earth (1978–2012), v. 110, no. B3. Parcheta, C., Houghton, B., and Swanson, D., 2013, Contrasting patterns of vesiculation in low, inter- mediate, and high Hawaiian fountains: A case study of the 1969 Mauna Ulu eruption: Journal of Volcanology and Geothermal Research. Parfitt, E., and Wilson, L., 1995, Explosive volcanic eruptions—IX. The transition between Hawaiian‐ style lava fountaining and Strombolian explosive activity: Geophysical Journal International, v. 121, no. 1, p. 226-232. Parfitt, E. A., 2004, A discussion of the mechanisms of explosive basaltic eruptions: Journal of Volcanol- ogy and Geothermal Research, v. 134, no. 1, p. 77-107. Park, J., Morgan, J. K., Zelt, C. A., and Okubo, P. G., 2009, Volcano-tectonic implications of 3-D velocity structures derived from joint active and passive source tomography of the island of Hawaii: Journal of Geophysical Research, v. 114, no. B9, p. B09301. Patrick, M., Orr, T., Swanson, D., and Lev, E., 2016a, Shallow and deep controls on lava lake surface motion at Kīlauea Volcano: Journal of Volcanology and Geothermal Research, v. 328, p. 247- 261. Patrick, M. R., Orr, T., Sutton, A. J., Lev, E., Thelen, W., and Fee, D., 2016b, Shallowly driven fluctuations in lava lake outgassing (gas pistoning), Kīlauea Volcano: Earth and Planetary Science Letters, v. 433, p. 326-338. Patrick, M. R., Orr, T., Wilson, D., Dow, D., and Freeman, R., 2011, Cyclic spattering, seismic tremor, and surface fluctuation within a perched lava channel, Kīlauea Volcano: Bulletin of volcanology, v. 73, no. 6, p. 639-653. Patrick, M. R., Swanson, D., and Orr, T., 2016c, Automated tracking of lava lake level using thermal images at Kīlauea Volcano, Hawai’i: Journal of Applied Volcanology, v. 5, no. 1, p. 1. Peters, N., Oppenheimer, C., Killingsworth, D. R., Frechette, J., and Kyle, P., 2014, Correlation of cycles in lava lake motion and degassing at Erebus volcano, Antarctica: Geochemistry, Geophys- ics, Geosystems, v. 15, no. 8, p. 3244-3257. Pietruszka, A. J., and Garcia, M. O., 1999, The size and shape of Kilauea Volcano’s summit magma storage reservoir: a geochemical probe: Earth and Planetary Science Letters, v. 167, no. 3, p. 311-320. Polacci, M., 2005, Constraining the dynamics of volcanic eruptions by characterization of pumice textures: Annals of Geophysics. Polacci, M., Burton, M. R., La Spina, A., Murè, F., Favretto, S., and Zanini, F., 2009, The role of syn-eruptive vesiculation on explosive basaltic activity at Mt. Etna, Italy: Journal of Volcanolo- gy and Geothermal Research, v. 179, no. 3, p. 265-269. Polacci, M., Corsaro, R. A., and Andronico, D., 2006, Coupled textural and compositional characterization of basaltic scoria: Insights into the transition from Strombolian to fire fountain activity at Mount Etna, Italy: Geology, v. 34, no. 3, p. 201-204. Poland, M., 2017, “Points Requiring Elucidation” About Hawaiian Volcanism, in Carey, R. J., Cayol, V., Poland, M., and Weis, D., eds., Hawaiian Volcanoes: From Source to Surface, Volume 208, John Wiley and Sons, Inc. Poland, M., Miklius, A., and Montogomery-Browm, E., A conceptual model of Kilauea’s magma plumbing system, in Proceedings AGU Fall Meeting, San Francisco, 2013. Porritt, L. A., Russell, J. K., and Quane, S. L., 2012, Pele’s tears and spheres: Examples from Kilauea 101

BibliographyIki: Earth and Planetary Science Letters, v. 333–334, p. 171-180. Reynolds, P., Brown, R. J., Thordarson, T., Llewellin, E. W., and Fielding, K., 2015, Rootless cone eruption processes informed by dissected tephra deposits and conduits: Bulletin of Volcanology, v. 77, no. 9, p. 1-17. Rhodes, J., 1988, Geochemistry of the 1984 Mauna Loa eruption: implications for magma storage and supply: Journal of Geophysical Research: Solid Earth, v. 93, no. B5, p. 4453-4466. Richter, D. H., Eaton, J., Murata, K., Ault, W., and Krivoy, H., 1970, Chronological narrative of the 1959-60 eruption of Kilauea volcano, Hawaii, 2330-7102. Rowan, L. R., and Clayton, R. W., 1993, The three‐dimensional structure of Kilauea Volcano, Hawaii, from travel time tomography: Journal of Geophysical Research: Solid Earth (1978–2012), v. 98, no. B3, p. 4355-4375. Rust, A., and Cashman, K., 2011, Permeability controls on expansion and size distributions of pyroclasts: Journal of Geophysical Research: Solid Earth, v. 116, no. B11. Ryan, M. P., 1988, The mechanics and three‐dimensional internal structure of active magmatic systems: Kilauea Volcano, Hawaii: Journal of Geophysical Research: Solid Earth (1978–2012), v. 93, no. B5, p. 4213-4248. Sable, J., Houghton, B., Wilson, C., and Carey, R., 2009, Eruption mechanisms during the climax of the Tarawera 1886 basaltic Plinian eruption inferred from microtextural characteristics of the deposits: Studies in Volcanology: The Legacy of George Walker, Spec. Publ. IAVCEI, v. 2, p. 129-154. Sable, J. E., Houghton, B. F., Del Carlo, P., and Coltelli, M., 2006, Changing conditions of magma ascent and fragmentation during the Etna 122 BC basaltic Plinian eruption: Evidence from clast microtextures: Journal of Volcanology and Geothermal Research, v. 158, no. 3, p. 333-354. Sahagian, D. L., and Proussevitch, A. A., 1998, 3D particle size distributions from 2D observations: stereology for natural applications: Journal of Volcanology and Geothermal Research, v. 84, no. 3, p. 173-196. Sawyer, G., Oppenheimer, C., Tsanev, V., and Yirgu, G., 2008, Magmatic degassing at Erta’Ale volcano, Ethiopia: Journal of Volcanology and Geothermal Research, v. 178, no. 4, p. 837-846. Sharp, W. D., and Clague, D. A., 2006, 50-Ma initiation of Hawaiian-Emperor bend records major change in Pacific plate motion: Science, v. 313, no. 5791, p. 1281-1284. Shea, T., Houghton, B. F., Gurioli, L., Cashman, K. V., Hammer, J. E., and Hobden, B. J., 2010, Textural studies of vesicles in volcanic rocks: an integrated methodology: Journal of Volcanology and Geothermal Research, v. 190, no. 3, p. 271-289. Smets, B., d’Oreye, N., Kervyn, M., and Kervyn, F., 2016, Gas piston activity of the Nyiragongo lava lake: First insights from a Stereographic Time-Lapse Camera system: Journal of African Earth Sciences. Spampinato, L., Ganci, G., Hernández, P., Calvo, D., Tedesco, D., Pérez, N., Calvari, S., Del Negro, C., and Yalire, M., 2013, Thermal insights into the dynamics of Nyiragongo lava lake from ground and satellite measurements: Journal of Geophysical Research: Solid Earth, v. 118, no. 11, p. 5771-5784. Spampinato, L., Oppenheimer, C., Calvari, S., Cannata, A., and Montalto, P., 2008, Lava lake surface characterization by thermal imaging: Erta’Ale volcano (Ethiopia): Geochemistry, Geophysics, Geosystems, v. 9, no. 12. Sparks, R., 1978, Effect of degassing on rheology of basaltic lava: Nature, .v 276, p. 385-386. Spilliaert, N., Allard, P., Métrich, N., and Sobolev, A., 2006, Melt inclusion record of the conditions of ascent, degassing, and extrusion of volatile-rich alkali basalt during the powerful 2002 flank eruption of Mount Etna (Italy): J. Geophys. Res, v. 111, p. B04203. 102

BibliographyStearns, H. T., and Macdonald, G. A., 1946, Geology and ground-water resources of the island of Ha- waii, Printed in the territory of Hawaii, v. 8-9. Stearns, H. T., Meinzer, O. E., and Clark, W. O., 1930, Geology and water resources of the Kau District, Hawaii. Stevenson, J. A., Mitchell, N. C., Mochrie, F., Cassidy, M., and Pinkerton, H., 2012, Lava penetrating water: the different behaviours of pāhoehoe and ‘a ‘ā at the Nesjahraun, Þingvellir, Iceland: Bulletin of volcanology, v. 74, no. 1, p. 33-46. Stovall, W. K., Houghton, B. F., Harris, A. J. L., and Swanson, D. A., 2009, A frozen record of density-driven crustal overturn in lava lakes: the example of Kīlauea Iki 1959: Bulletin of Volcanol- ogy, v. 71, no. 3, p. 313-318. Stovall, W. K., Houghton, B., Gonnermann, H., Fagents, S., and Swanson, D., 2011, Eruption dynamics of Hawaiian-style fountains: the case study of episode 1 of the Kīlauea Iki 1959 eruption: Bul- letin of volcanology, v. 73, no. 5, p. 511-529. Stovall, W. K., Houghton, B. F., Hammer, J. E., Fagents, S. A., and Swanson, D. A., 2012, Vesiculation of high fountaining Hawaiian eruptions: episodes 15 and 16 of 1959 Kīlauea Iki: Bulletin of volcanology, v. 74, no. 2, p. 441-455. Swanson, D. A., Duffield, W. A., and Fiske, R. S., 1976, Displacement of the South Flank of Kilauea Volcano: The Result of Forceful Intrusion of Magma Into the Rift Zones: Interpretation of Geo- detic and Geologic Information Leads to a New Model for the Structure of Kilavea Volcano, US Government Printing Office. Swanson, D. A., Duffield, W. A., Jackson, D. B., and Peterson, D. W., 1979, Chronological narrative of the 1969-71 Mauna Ulu eruption of Kilauea volcano, Hawaii: US Govt. Print. Off., 2330-7102. Swanson, D. A., and Fabbi, B. P., 1973, Loss of volatiles during fountaining and flowage of basaltic lava at Kilauea volcano, Hawaii: Journal of Research of the U. S. Geological Survey, v. 1, no. 5, p. 649-658. Swanson, D. A., Rose, T. R., Mucek, A. E., Garcia, M. O., Fiske, R. S., and Mastin, L. G., 2014, Cycles of explosive and effusive eruptions at Kīlauea Volcano, Hawai ‘i: Geology, v. 42, no. 7, p. 631- 634. Taddeucci, J., Edmonds, M., Houghton, B. F., James, M., and Vergniolle, S., 2015, Hawaiian and Strom- bolian Eruptions, in Sigurdsson, H., ed., Encyclopedia of Volcanoes, Elsevier. Tazieff, H., 1994, Permanent lava lakes: observed facts and induced mechanisms: Journal of Volcanolo- gy and Geothermal Research, v. 63, no. 1, p. 3-11. Thorarinsson, S., 1951, Laxárgljúfur and Laxárhraun: a tephrochronological study: Geografiska annaler, p. 1-89. Thorarinsson, S., 1951, 1953, The crater groups in Iceland: Bulletin of volcanology, v. 14, no. 1, p. 3-44. Thordarson, T., Rootless eruptions and cone groups in Iceland: Products of authentic explosive water to magma interactions, in Proceedings Volcano–Ice Interactions on Earth and Mars Conference. Reykjavık2000, p. 48. Thordarson, T., and Höskuldsson, Á., 2008, Postglacial volcanism in Iceland: Jökull, v. 58, p. 197-228. Thordarson, T., Miller, D., and Larsen, G., 1998, New data on the age and origin of the Leidólfsfell cone group in south Iceland: Jökull, v. 46, p. 3-15. Toramaru, A., 1995, Numerical study of nucleation and growth of bubbles in viscous magmas: Journal of Geophysical Research: Solid Earth, v. 100, no. B2, p. 1913-1931. Valentine, G. A., and Connor, C. B., 2015, Basaltic volcanic fields, The Encyclopedia of Volcanoes (Second Edition), Elsevier, p. 423-439. Vergniolle, S., and Gaudemer, Y., 2015, From reservoirs and conduits to the surface: Review of role of 103

Bibliographybubbles in driving basaltic eruptions: Am. Geophysical Union Monograph:“Hawaiian Volca- nism: From Source to Surface, v. 208, p. 289-322. Vergniolle, S., and Jaupart, C., 1990, Dynamics of degassing at Kilauea volcano, Hawaii: Journal of Geophysical Research, v. 95, no. B3, p. 2793-2809. Vergniolle, S., and Bouche, E., 2016, Gas-driven lava lake fluctuations at Erta’Ale volcano (Ethiopia) revealed by MODIS measurements: Bulletin of Volcanology, v. 78, no. 9, p. 60 Walker, G., and Croasdale, R., 1971, Characteristics of some basaltic pyroclastics: Bulletin Volca- nologique, v. 35, no. 2, p. 303-317. Walker, G. P., 1973, Explosive volcanic eruptions—a new classification scheme: Geologische Rund- schau, v. 62, no. 2, p. 431-446. Walker, G. P., 1992, Puu Mahana Near South Point in Hawaii Is a Primary Surtseyan Ash Ring, Not a Sandhills-type Littoral Cone. Walker, G. P., 1993, Basaltic-volcano systems: Geological Society, London, Special Publications, v. 76, no. 1, p. 3-38. Walter, T. R., and Amelung, F., 2006, Volcano‐earthquake interaction at Mauna Loa volcano, Hawaii: Journal of Geophysical Research: Solid Earth, v. 111, no. B5. White, J. D., Smellie, J. L., and Clague, D. A., 2003, Explosive subaqueous volcanism, American Geo- physical Union Washington, DC. Wilson, D., Elias, T., Orr, T., Patrick, M., Sutton, J., and Swanson, D., 2008, Small explosion from new vent at Kilauea’s summit: Eos, Transactions American Geophysical Union, v. 89, no. 22, p. 203. Wilson, L., 1980, Relationships between pressure, volatile content and ejecta velocity in three types of volcanic explosion: Journal of Volcanology and Geothermal Research, v. 8, no. 2, p. 297-313. Wilson, L., and Head, J., 1981, Ascent and eruption of basaltic magma on the Earth and Moon: Journal of Geophysical Research, v. 86, no. B4, p. 2971-3001. Witham, F., and Llewellin, E. W., 2006, Stability of lava lakes: Journal of Volcanology and Geothermal Research, v. 158, no. 3–4, p. 321-332. Wohletz, K., 2002, Water/magma interaction: some theory and experiments on peperite formation: Jour- nal of Volcanology and Geothermal Research, v. 114, no. 1, p. 19-35. Wohletz, K. H., 1983, Mechanisms of hydrovolcanic pyroclast formation: grain-size, scanning electron microscopy, and experimental studies: Journal of Volcanology and Geothermal Research, v. 17, no. 1, p. 31-63. Wolfe, C. J., Brooks, B. A., Foster, J. H., and Okubo, P. G., 2007, Microearthquake streaks and seismicity triggered by slow earthquakes on the mobile south flank of Kilauea Volcano, Hawai'i: Geophysical Research Letters, v. 34, no. 23. Wolfe, E., Garcia, M., Jackson, D., Koyanagi, R., Neal, C., and Okamura, A., 1987, The Puu O eruption of Kilauea Volcano, Episodes 1-20, January 3, 1983, to June 8, 1984, in Decker, R., Wright, T., and Stauffer, P., eds., Volcanism in Hawaii, Volume 1, U.S. Government Printing Office. Wolfe, E., Neal, C., Banks, N., and Duggan, T., 1988, Geologic observations and chronology of eruptive events: US Geol Surv Prof Pap, v. 1463, p. 1-97. Wolfe, E. W. M., Jean, 1996, Gelogic Map of the Island of Hawaii, Map I-2524-A, U.S. Geological Survey. Wright, T. L., and Klein, F. W., 2006, Deep magma transport at Kilauea volcano, Hawaii: Lithos, v. 87, no. 1, p. 50-79. Zimanowski, B., Büttner, R., Dellino, P., White, J. D., and Wohletz, K. H., 2015, Magma–Water Inter- action and Phreatomagmatic Fragmentation, The Encyclopedia of Volcanoes (Second Edition), Elsevier, p. 473-484. 104

Zimanowski,Bibliography B., Büttner, R., and Lorenz, V., 1997, Premixing of magma and water in MFCI experiments: Bulletin of volcanology, v. 58, no. 6, p. 491-495. Zurek, J., and Williams-Jones, G., 2013, The shallow structure of Kilauea caldera from high-resolution Bouguer gravity and total magnetic anomaly mapping: Insights into progressive magma reservoir growth: Journal of Geophysical Research: Solid Earth, v. 118, no. B7, p. 3742-3752.