PRPF4 is a novel therapeutic target for the treatment of breast cancer by influencing growth, migration, invasion, and apoptosis of breast cancer cells via p38 MAPK signaling pathway

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Park Jeffrey, John (Orcid ID: 0000-0001-8776-270X)

Why is Crustal Underplating Beneath Many Hotspot Islands Anisotropic? Jeffrey Park1 and Danny M. Rye1

1

Dept. of Geology & Geophysics, Yale University, New Haven, Connecticut, USA

Submitted to G-Cubed, June 05, 2019 Corresponding author: Jeffrey Park ([email protected]) Key Points: 

Active-source seismics and receiver functions find that many hotspot islands have underplated layers, with Vp intermediate to crust and mantle



High-frequency receiver functions from two GSN stations in the Pacific exhibit complex anisotropic layering, extending a study at lower frequency



We propose that seawater descends thermal cracks and serpentinizes sub-Moho mantle, lowering density, Vp, and adding anisotropy via crack textures

Abstract

High-frequency harmonic regression (2.0-4.0-Hz cutoff) of receiver functions at two long-

running seismic observatories at mid-Pacific hotspot islands confirms earlier detections of this underplated material with seismic velocities intermediate to crust and mantle, and reveals it to be multilayered and anisotropic within ~30 km of the surface. Magmatic underplating beneath

the oceanic Moho has been proposed to accompany basaltic melt that erupts at the seafloor and (eventually) atop a subaerial volcano. An alternate hypothesis is “metasomatic underplating”

whereby crustal fractures developed during magma ascent allow seawater to infiltrate and to serpentinize the sub-Moho mantle partially. Metasomatic underplating would lower seismic wavespeeds, promote the buoyancy of the hotspot swell, and induce textural anisotropy as metamorphic expansion of olivine-rich peridotite promotes a crack network along which serpentinization spreads. Differential expansion of mantle peridotite and crustal gabbro This article has been accepted for publication and undergone full peer review but has not been through the copyediting, typesetting, pagination and proofreading process which may lead to differences between this version and the Version of Record. Please cite this article as doi: 10.1029/2019GC008492 ©2019 American Geophysical Union. All rights reserved.

promotes cracks in the crust that offer new pathways for seawater to descend to the Moho,

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allowing metasomatic underplating to expand laterally and to contribute anisotropy to the

underplated layer. Rare serpentinized mantle xenoliths confirm that crack textures can develop during serpentinization at depth. The discovery of iron-oxidizing microbial mats on the seafloor flank of the Loihi volcano, and many locations of diffuse low-temperature venting worldwide, is consistent with the circulation of metasomatic fluids with reducing chemistry, sourced from serpentinization at depth. Post-eruptive uplift of Santa Maria Island (Azores), and asymmetry of the Hawaiian swell, suggests that underplating requires 2-4 Myr to complete, suggesting that fluid infiltration is slow, subject to cycles of blockage and fresh fracturing. Plain Language Summary: A 5-10-km layer of rock of intermediate seismic velocities

and strong anisotropy is often found beneath volcanic hotspot islands, lying between the oceanic crust and mantle. This underplated layer has typically been assumed to form as a reservoir of magma that fails to erupt at the seafloor or on the island, but this model predicts several processes and features, such as shallow seismic activity and a ring of raised topography, that are matched by parallel geologic processes at the Yellowstone hotspot and in the coronae of Venus. We propose a model in which seawater infiltrates the oceanic crust to convert mantle

rock partially to the mineral serpentine, which lowers both seismic velocities and densities of the rock layer, as well as generating elastic anisotropy via crack networks that develop as the chemical reactions proceed. The underplated layer adds buoyancy to the swells that surround hotspot islands. Using the recent dating of an uplifted island in the Azores, we estimate that 2-4 million years are necessary for this underplated layer to complete its growth. Index terms: 7200 Seismology, 3000 Marine Geology and Geophysics, 8100 Tectonophysics, 1000 Geochemistry, 5400 Planetary Sciences: Solid Surface Planets Keywords: hotspot, mantle plume, oceanic plate, seismic anisotropy, receiver functions, serpentinite, Hawaii, metasomatism, underplating

1: Introduction

Active-source marine refraction and reflection seismology deployments [e.g., ten Brink and

Brocher, 1987; Wolfe et al., 1994; Caress et al., 1995; Gallart et al., 1999; Grevemeyer et al., 2001; Canales et al., 2002; Contreras-Reyes et al., 2010] across diverse volcanic-island systems find evidence for a layer beneath the oceanic crust, often interpreted as magmatic underplating by plutonic rock of intermediate wavespeeds. A distinct layer beneath the oceanic Moho has

been reported also from receiver-function (RF) studies [Leahy and Park, 2005; Lodge and Helffrich, 2006; Lodge et al., 2012; Rychert et al., 2014], suggesting magmatic underplating with compositional buoyancy sufficient to explain much of the hotspot swell [Lodge and

©2019 American Geophysical Union. All rights reserved.

Helffrich, 2006; Leahy et al., 2010]. In the case of Hawaii, data from ocean-bottom

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seismometers has detected a sub-Moho layer, attributed to magmatic underplating, that extends 100s of km from the loci of active hotspot volcanism [Leahy et al., 2010]. Olugboji and Park [2016] reported that the underplated layer at 11 hotspot islands contained anisotropy that weakly correlated with inferred stress directions at the time of island formation. In this paper we refer to “underplating” as a structural feature, irrespective of origin.

“Magmatic underplating” is the common understanding of the term, implying a plutonic origin. “Diapiric underplating” could be effected by a low-viscosity solid, perhaps a thermal plume containing partial melt. A third option is “metasomatic underplating,” in which a buried layer

of rock, typically peridotite, is infiltrated by water to form a hydrated rock of lowered density, primarily via the production of serpentine [e.g., Bach et al., 2006]. We note that peridotite hydration via the formation of chlorite [Mookherjee and Mainprice, 2014] or lawsonite [Broverone and Beyssac, 2014] may occur in tandem with serpentinization, depending on local conditions. Crack-mediated fluid flow during serpentinization is a likely source of rock texture that leads to elastic anisotropy in the underplated layer. Is ocean-island underplating caused by radial spreading of hot mantle rock or magma from

the plume conduit to form a tabular sill? Lateral spreading of a shallow low-viscosity droplet near the free surface was modeled by Koch and Koch [1995]. This geodynamic model generates concentric rings of radial extension and compression at the free surface above the edge of the spreading body. Topographic rings developed from this stress pattern were proposed to explain the coronae of Venus [Koch, 1994; Koch and Manga, 1996]. No such

bathymetric features have been recognized at the edges of Earth’s mid-ocean hotspot swells. This paper outlines an alternate hypothesis for underplating beneath hotspot islands in

which a low-velocity layer develops beneath oceanic crust via serpentinization of ultramafic rock just beneath the Moho. When ultramafic rocks serpentinize, they expand in volume. At seafloor exposures near the slow-spreading mid-Atlantic ridge, this metamorphic volume change helps to form the updomed Atlantis massif at 30°N [Schroeder et al., 2002] and the Lost City hydrothermal field [Kelley et al., 2005]. We propose that the onset of hotspot volcanism creates crustal fractures that penetrate mature ocean crust, allowing seawater to circulate downward to serpentinize ultramafic rock below the Moho. Section 2 reviews evidence that volcanic mid-ocean islands are underplated by low

wavespeed rock, and that the underplated layer is anisotropic. We focus on two permanent seismological observatories that have been discussed in previous studies. Leahy and Park [2005] estimated radial receiver functions with frequency cutoffs up to fC=4.0Hz to detect thin ©2019 American Geophysical Union. All rights reserved.

layered structures. Olugboji and Park [2016] estimated both radial and transverse RFs with

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fC=1.0Hz with harmonic regression [Bianchi et al., 2010; Schulte-Pelkum and Mahan, 2014; Park and Levin, 2016b] to resolve anisotropy within the sub-Moho underplated layer. Here we apply the harmonic RF regression of the latter study with the higher frequency cutoffs of Leahy and Park [2005] to provide a high-resolution sample of seismic structures that must be explained by any hotspot-underplating model, and to reconcile previous analyses. In the following sections we present evidence from diverse subdisciplines that a

serpentinization scenario can be taken seriously as an explanation for underplating beneath mid-ocean islands. We cannot verify the hypothesis completely with currently available data, but several avenues for further investigation are apparent. Section 3 discusses the pressure and temperature constraints associated with serpentinization below the oceanic Moho. Section 4 discusses geophysical evidence, particular to suggest how our serpentinization hypothesis could address the venerable oceanic heat-flow paradox [Stein and Stein, 2015], and unexpected patterns of seismicity near Hawaii [Klein, 2016]. Section 5 discusses supporting evidence from diffuse fluid-vents and Zetaproteobacteria microbial communities near Hawaii and elsewhere [Emerson and Moyer, 2010], the potential implications for our understanding of hotspot-lava geochemistry [Wang and Eiler, 2008], cycling of water in Earth’s mantle [Kendrick et al., 2017], flat slabs [Manea et al., 2017], displaced arc volcanoes [Kay and Mpodozis 2002; Nikulin et al., 2010], post-emplacement uplift of volcanic islands [Ramalho et al., 2015; 2017], and the geodynamical metrics of hot-spot swells [Morgan et al., 1995; Harris and McNutt, 2007; King and Adam, 2014]. Section 6 summarizes the case for the metasomatic underplating (MSUP) hypothesis. 2: Seismic Evidence for Anisotropic Underplated Layers Active-source seismic surveys and receiver functions have found that many hotspot islands

have underplated layers, with seismic velocities intermediate to crust and mantle. The structural details of an underplated layer are often not the primary focus of a seismological experiment. Evidence collected over several decades has not clarified the origin of these features. In this section we increase the frequency content of receiver functions estimated from two permanent ocean-island stations to resolve fine-layered 1-D anisotropic structure. Multilayered anisotropy, which develops within relatively rigid oceanic lithosphere, is an important clue to the formation mechanism of the underplated layer. 2.1: Previous Studies

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In the case of Hawaii, the sub-Moho underplated layer extends far beyond the footprint of

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the volcanic edifices, posing problems of its geodynamic origin. An active-source seismic survey [ten Brink and Brocher, 1987] detected a 3-6-km thick layer up to 100 km north and

south of Oahu in the Hawaiian Islands. They proposed a spreading magmatic intrusion, fed from a magmatic source centered on Oahu, that created a plutonic body with an aspect ratio of 30:1 or more. Receiver-functions from ocean-bottom seismometer (OBS) data collected on the Hawaiian plume swell [Leahy et al., 2010] challenge a spreading-plume model for the underplated layer. OBS RFs display evidence for an underplated layer as distant as 200 km from the active loci of volcanism, but no underplating was detected beneath a salient of seafloor north of Hawaii that intersects the Molokai Fracture Zone -- see Figure 12 of Leahy et al. [2010] for the geographic pattern and ten Brink and Brocher [1988] for confirmation that no underplating is evident beneath the Molokai Fracture zone. The gap in the underplated layer is inconsistent with the lateral spreading of a magmatic or low-viscosity sill. Leahy et al. [2010] proposed that basaltic magma has intruded the underplated layer from below, relating gaps in the underplated layer to unknown complexities in the Hawaiian plume head. A recent study of the Hawaii plume head with surface-wave dispersion estimated from

seismic data from the same seafloor deployment [Laske et al., 2011] does not address the underplated layer. It confines its largest VS perturbations to the 80-140-km depth range, beneath the nominal depth of the lithosphere-asthenosphere boundary in mature ocean [Kawakatsu et al., 2009; Rychert and Shearer, 2011; Olugboji et al., 2016], though it also infers smaller velocity anomalies within the lithosphere. Lateral variations in VS are evident in the Laske et al. [2011] study, but not all correlate with the reports of underplating by Leahy et al. [2010]. In one example of agreement, at 40-km depth Laske et al. [2011] infers a localized patch of weak or absent VS anomaly beneath the seafloor patch NW of Hawaii and Maui where underplating is not seen. However, Laske et al. [2011] report a strong VS deficit at 40km depth south of Hawaii beneath the South Arch Volcanic Field, where underplating is also not seen. The first comparison supports the causal hypothesis of Leahy et al. [2010] for the underplated layer, that is, magmatic infiltration from below, but the second comparison does not. P-wave tomography [Wolfe et al., 2011] and Sp receiver functions [Rychert et al., 2013]

lack sufficient resolution in the crust and upper 50 km of the mantle to resolve structural detail relevant to the underplated layer. Receiver-function studies are easy to replicate and to improve by the addition of more

data over time. Leahy and Park [2005] identified an underplated layer beneath each of four

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permanent ocean-island seismological observatories (RAR, PPT, XMAS, POHA).

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Identification of thin layers was aided by applying frequency cutoffs of 2.0-4.0 Hz, with bandpass-tapering in the frequency domain to suppress ringing [Park and Levin, 2000; 2016a]. Distinct converted Ps phases could often be identified from an interface near the base of the volcanic edifice, the original oceanic Moho, and the bottom of an underplated layer beneath the Moho. Leahy et al. [2010] revised the interpretation of the shallowest interface, at least in Hawaii, to be the transition from subaerial to subaqueous material in the extrusive volcano. Other researchers have also inferred crustal underplating from receiver functions, but their interpretations of layer thicknesses have been influenced by the frequency content of

their analyses. Lodge and Helffrich [2006], Lodge et al. [2012], Fontaine et al. [2015], Liu and Park [2017] and Spieker et al. [2018] applied frequency cutoffs sufficiently high to resolve short pulses. They report underplated layers beneath the Canary Islands, La Reunion,

Hawaii and Sao Jorge in the Azores with thicknesses of 10 km or less. Rychert et al. [2014] estimate Sp receiver functions, which typically involve lower frequencies than Ps receiver functions, to infer an underplated layer beneath the Galapagos Islands that extends to 37±7 km depth. Olugboji and Park [2016] reported underplated layers beneath 11 ocean-island stations within the Pacific Ocean basin, using a harmonic regression of Ps receiver functions over back azimuth to investigate anisotropic layering with frequency cutoff 1 Hz. With this lowpassed restriction, Olugboji and Park [2016] typically inferred thick layers beneath their

stations, roughly 30 km for the combination of crust and underplated layer. Olugboji and Park [2016] investigated elastic anisotropy beneath 11 permanent seismic

stations on Pacific hotspot islands, including two Hawaii stations, by estimating back-azimuth harmonics of RFs. They found a two-lobed variation in Ps converted-wave amplitude whose minimum typically aligns within 30° of plate motion at the time of volcano emplacement. The RFs were consistent with anisotropy with a “slow” tilted symmetry axis localized within the underplated layer. Because the underplated layer lies within the stiff lithosphere, alignment with plate motion cannot be assumed from asthenospheric drag. Olugboji and Park [2016] proposed a model in which upward migration of basaltic magma rises though lithospheric fractures tilted by stresses induced by the local bending of lithosphere from edifice loading [Hieronymus and Bercovici, 1999; 2000], see also Richards et al. [2013]. Aside from tilt, the induced or reactivated fractures would orient roughly with the extensional stresses that drag the plate toward a subduction zone, explaining the slow-axis orientations inferred within the underplated layer. This model shares with Leahy et al. [2010] the hypothesis that ocean-island underplating is determined by vertical, rather than lateral, migration of low-viscosity material, ©2019 American Geophysical Union. All rights reserved.

in this case infiltrating melts. Local-bending stresses to channel basaltic melts may not be

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sufficiently broad, however, to open conduits to supply a large plume swell. 2.2: Receiver-Functions at Two Mid-Ocean Islands at High Frequency

In this paper we revisit two stations analyzed by both Leahy and Park [2005] and

Olugboji and Park [2016], applying the back-azimuth harmonic regression of the latter study at a high-frequency cutoff similar to the former study. We present these analyses in considerable detail, because the harmonic-regression technique is somewhat new to the receiver-function community. We apply trial-and-error forward-modelling of receiverfunctions with synthetic seismograms from reflectivity simulations of wave propagation in simple 1-D structures, using the code described by Levin and Park [1997]. One pitfall of this approach is that synthetic RFs often contain layer reverberations [e.g., Zhu and Kanamori, 2000] that may not be observed in the data-derived RFs. A layer reverberation requires a flat, sharp interface to coalesce converted-wave energy into coherent pulses after multiple reflections. Interfaces of such simplicity are often lacking in the neighborhood of a midocean volcanic island. As a result, one must often decide whether a prominent reverberation phase in a synthetic RF should be ignored when assessing the data-fit of a modeling experiment. This issue arises for both stations we report here. We also use updated data sets. Our targets are two representative stations of the Global

Seismographic Network, POHA (Pohakuloa, Hawaii) and RAR (Rarotonga Island) [Butler et al., 2004; Albuquerque Seismological Laboratory, 1988]. Liu and Park [2017] report RFs for a third hotspot-island station (KIP, Kipapa, Hawaii) [Albuquerque Seismological Laboratory, 1988] that confirms the general observations that we make in this study. Our goals include

(1) confirming the existence of a distinct underplated layer beneath the nominal oceanic Moho, (2) getting a rough estimate of seismic velocities and anisotropy in the underplated layer, and (3) determining whether the underplated layer has uniform, gradational, or multilayered anisotropic properties. Considered in isolation, the multi-layered structure of seismic properties beneath mid-ocean islands will not uniquely favor one causal hypothesis for the underplated layer over another. However, the layering style will influence possible scenarios. For example, if high resolution RFs confirm an interpretation in terms of simple 510-km thick layers with uniform properties, simple one-step causal processes, such as a massive magmatic intrusion, are more likely. Multi-layered structure might favor a multistep causal process, in which perturbations to oceanic lithosphere increment over time. We estimate Ps receiver functions from multiple-taper correlation (MTC) as described by

Park and Levin [2000], using T=60-s time windows that are Fourier-transformed with three ©2019 American Geophysical Union. All rights reserved.

Slepian tapers with time-bandwidth product p=2.5. We apply LQT rotation before spectrum

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analysis [Reading et al., 2003], aligning the vertical and radial-horizontal components of the seismograms parallel and perpendicular, respectively, to the predicted polarization of the P wave. We use moveout delays in the Q and T motion components to target specific depth ranges for the RFs [Helffrich, 2006; Bianchi et al., 2010]. Our discussion will typically defer to common vernacular and refer to the longitudinal (L) component as “vertical” and the Q component as “radial,” because the moveout corrections we apply shift all delay times to those appropriate for vertical-incidence P waves [Park and Levin, 2016a]. The Slepian tapers lead to spectrum estimates that average information over a total bandwidth f=5fR, where fR=1/T=0.0167Hz is the Rayleigh frequency of the 60-s data windows. We compute receiver functions via linear regression rather than spectral division, thereby obviating water-level damping that can drown high-frequency correlations in the seismic spectra. Anisotropy with a vertical symmetry axis generates no P-SH conversion for a horizontal

interface, and generates a P-SV amplitude that is constant with back azimuth. A horizontal symmetry axis leads to a four-lobed Ps pattern, varying with cos(2(P-SV) or sin(2(PSH), where  is the counter-clockwise angle between back-azimuth and the symmetry-axis strike [Booth and Crampin, 1983; Kosarev et al., 1984; Savage, 1998; Levin and Park, 1998a;

Eckhardt and Rabbel, 2011]. A dipping interface or a symmetry axis that tilts at an angle 

from the vertical adds a two-lobed Ps pattern, varying with cos (P-SV) or sin (P-SH). The relative phasing of P-SV and P-SH converted-wave amplitudes with back azimuth can be

stacked for both 2-lobed and 4-lobed components [Park and Levin, 2016b]. The backazimuth variations of P-SV conversion are 90° phase-advanced relative to P-SH, and this phase combination can be modeled by anisotropy and/or dipping interfaces [Schulte-Pelkum and Mahan, 2014]. In the opposite phasing P-SV conversions are 90° phase-lagged relative

to P-SH in back azimuth. We label this phase-combination “Unmodeled.” The Unmodeled

amplitude indicates deviation from assumptions of flat-layered anisotropy in the data set. 2.2: Fine-Layering Without Reverberation: Station RAR Volcanic activity at Rarotonga, Cook Islands was active during 1.1-2.5 Ma [Thompson et

al., 1998]. GSN station RAR is sited in volcanic rock 28 meters above sea level. We accessed borehole-sensor data on location code 00 for 1999-2016 events (either a KS-36000

or KS-54000 sensor), and surface-vault data on location code 10 from 2003-2016 (either a CMG-3T or STS-2 sensor). Recurrent data-quality issues limited us to select 246 and 206 events for location codes 00 and 10, respectively, for a total of 302 unique events (Figure 1).

©2019 American Geophysical Union. All rights reserved.

Stacked receiver functions computed separately for each location code agreed well, so we

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combined the data sets for a composite estimate. We spline-interpolated the 20-Hz-sampled data from location code 00 to 40 Hz sampling to match the data from location code 10. Olugboji and Park [2016] found RAR to be a problematic station, with mis-reported horizontal-component alignments. After attempts at correction, Olugboji and Park [2016] reported harmonic RF stacks with larger uncertainties than all other stations in their data set, except JOHN (Johnston Atoll, USA). For this study, we made fresh requests of seismic data from RAR from the IRIS Data Management System to exploit updates to its metadata. Our data-quality criteria were stringent, so that our dataset included ~100 fewer events than used by Olugboji and Park [2016]. The results of re-analysis were encouraging, with small bootstrap uncertainties in the harmonic RF regressions and modest “unmodeled” amplitude. An epicentral sweep of RFs, using events indicated in Figure 1, illustrates the polarity and

moveout of major Ps converted phases. In the synthetic RF sweep (Figure 2), the direct Ps phase from the deepest interface (17 km) is obscured by the reverberations of the shallowest interface (4 km), and is evident only at the smallest epicentral distances. For the RAR data, the characteristic positive-negative polarity pairing of the Ppms and Psms reverberations from the shallowest interface is absent, leaving a single positive pulse that we identify with a direct Ps conversion. In another comparison, a negative-polarity pulse in the data-derived RF sweep at 8-10-s time delay follows the moveout predicted for the Psms reverberation of the 17-km stacking-model interface, but no positive-polarity Ppms reverberation is evident. A similar disagreement applies to the predicted reverberations of the 9-km interface, corresponding to the original oceanic Moho. Receiver-functions from seafloor stations can be afflicted by water-column reverberations

[Bostock and Trehu, 2012; Janiszewski and Abers, 2015], but the bathymetry surrounding Rarotonga Island does not seem favorable for trapped-mode energy. Summerhayes [1967] describes Rarotonga as a volcanic cone with a 50-km diameter base and concave slopes that increase from 2° at 4-km seafloor depth to 20°-30° in shallow waters. Rarotonga has fringing reefs, but it is not an atoll, so it lacks a lagoon. The seafloor geometry does not seem to favor water-reverberations at discrete repeat-times. Comparing data against synthetic RFs, we conclude that a negative direct-Ps phase is evident near 7-s delay time, diagnostic of a velocity drop with depth. One may also infer a fourth positive direct-Ps pulse near 3-sec delay, suggesting a second underplated layer. We computed harmonic regressions of the event RFs (Figure 3) using moving-window

migration [Helffrich, 2006; Park and Levin, 2016a] and bootstrap resampling to estimate the ©2019 American Geophysical Union. All rights reserved.

uncertainty of the time traces. The stacking model for moveout corrections can be found in

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Table 1. Harmonic regression assumes that the Ps amplitude varies with earthquake backazimuth with a linear combination of cosine and sine functions, which form the “harmonics,” and fits for the expansion coefficients via weighted least-squares. The radial and transverse RF amplitudes are predicted to vary in quadrature in the presence of anisotropy and/or dipping interfaces [Schulte-Pelkum and Mahan, 2014; Licciardi and Agostinetti, 2016; Park and Levin, 2016b], so we regress for a composite radial/transverse RF. The radial/transverse combination with the opposite sense of quadrature is termed “Unmodelled;” its amplitude can

be used to assess whether the modeling assumptions are robust. The relative size of the “Anisotropy/Dip” and “Unmodelled” traces confirms this robustness. Even the low-amplitude cos(2) and sin(2) terms exceed their “Unmodelled” combinations by 2X for most delay times. Olugboji and Park [2016] reported substantial “Unmodelled” amplitude for RAR, but analysis of our dataset does not. Our selection and processing of body-wave records was

independent of the earlier study, and encountered none of the horizontal-component misalignment issues reported by Olugboji and Park [2016]. In the “constant” RF trace, a smaller, but substantial, positive pulse at 4-s time delay

supports the suggestion from Figure 2 for a second sub-Moho layer. (This converted phase is evident, but weaker, in the narrow back-azimuth sector plotted in Figure 2.) Leahy and Park [2005] identified the 4-s-delay phase at RAR with a Ppms reverberation within the oceanic crust (exclusive of the volcanic layer), but the moveout of this phase is opposite to that of a

reverberation (Figure 2, see purple line at 4.0-4.3-sec time delay). The negative pulses at 510-s time delay could include some reverberation signal, but Figure 2 shows that interpretation to be problematic. Upon close inspection, the two purple lines in Figure 2 at 46-sec delay time mark reverberations that should have positive (4.0-4.3 sec) and negative (5.3-5.8 sec) polarities, but do not, and the same can be seen for the possible reverberations marked by green lines at 6.5-7.2-sec and 9.0-9.6-sec delay times. Because the polarity test for reverberations fails, we infer at least one seismic-wavespeed inversion in the 40-80-km depth interval beneath RAR, and probably more than one. These interfaces likely are not horizontal, because the epicentral moveout of the negative Ps phases in Figure 2 is not simple.

RF variation that is 2-lobed in back azimuth (cos() and sin() terms) has far larger

amplitude than 4-lobed variation, indicating the presence of substantial tilted-axis anisotropy (Figure 3). The largest signals occur at delay times consistent with Ps conversions at the top

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and bottom of the oceanic crust, but substantial signals are also evident in the 2-7-s interval.

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The orientation of Ps maximum amplitudes and polarity flips are not consistent with the anisotropic symmetry-axis orientation (305°) reported by Olugboji and Park [2016]. The substantial changes in RF properties are likely associated with improved metadata for RAR.

In = the Supplemental Information, Figure S2 shows estimates of radial and transverse RFs in narrow sectors of back azimuth, estimated for cutoff frequencies fC =1.0 Hz, 2.0 Hz, 4.0 Hz and 6.0 Hz. In these back-azimuth plots the shallowest (time delay t < 1s) transversecomponent 2-lobed patterns of Ps converted phases flip polarity near 240° and 60° back azimuth (clockwise from north), while those at ~2.5-sec delay time express opposite polarity

on the transverse RFs with peak-amplitude near 80°-160° and 265°-330° back azimuth. Both these variations suggest a symmetry axis that strikes roughly NE-SW, consistent with the harmonic-regression RF estimates. We applied moveout corrections [Park and Levin, 2016a] to align the receiver functions

at 30-km target depth, near the Ps at the deepest potential depth of the underplated layer beneath RAR (Figure 4). We applied a different set of moveout corrections to target 70-km depth, roughly the depth of the seismic lithosphere-asthenosphere boundary (LAB) beneath old ocean floor [Kawakatsu et al., 2009; Rychert and Shearer, 2011; Olugboji et al., 2016]. As the target depth deepens, Ps signals from shallow interfaces defocus and Ps signals from deeper interfaces sharpen, as long as the deeper interfaces are well-approximated by horizontal planes [Liu et al., 2018]. For the 30-km depth target (Figure 4a) the Anisotropy/Dip harmonic terms are prominent. Ps pulses on the cos() and sin() traces are more closely-spaced than on the trace that is constant with back azimuth, suggest that layering in the anisotropy is finer than layering in isotropic seismic velocities. For the 70-km depth target (Figure 4b) the constant harmonic term displays a broad negative-polarity feature that suggests a gradual velocity inversion well above the expected depth of the seismic LAB. Because the Rarotonga geoid high is modest [Lambeck and Coleman, 1982], it is unlikely that thermal erosion has thinned the lithosphere by 20 km or more. In the narrow backazimuth section plotted in Figure 2, the negative Ps signals at 5-9-s time delay are not broad,

but rather manifest as multiple negative Ps pulses diagnostic of separate interfaces, see also

Leahy and Park [2005]. It is likely that this deeper feature departs from 1-D structure. Olugboji and Park [2016] argued that anisotropy alignment beneath station RAR

supported a hypothesis of fracture formation in the sub-Moho mantle, with strikes normal to plate-motion extension at the time of emplacement. In our analysis with better constraints of the horizontal-component orientation, RAR no longer aligns well with past plate motion. ©2019 American Geophysical Union. All rights reserved.

Moreover, with the higher RF frequency cutoff that we report here, substantial thin-layered

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anisotropy in crustal layers is evident, with significant variations in geometry. The crustal orientation (60° CW from N) does not align with the roughly north-south paleo-spreading direction of this portion of the SW Pacific [Williams et al., 2016; Boschman and van Hinsbergen, 2016], so fossil fractures do not explain it. Figure 5 shows the “Anisotropy/Dip” harmonic components of RFs estimated from synthetic P-coda computed for the model in Table 2, obtained with trial-and-error forward modelling. Key features of these synthetics include layer reverberations, particularly a doublet of Ppms/Psms phases in the first few seconds from the shallow volcanic layer. The shallow reverberative phases in the synthetics have no counterparts in the RFs estimated from RAR data. Park and Levin [2016b] demonstrate that the influence of S-anisotropy on Ps scattering

for tilted-axis anisotropy is small, so we fixed its amplitude relative to P-anisotropy. SubMoho anisotropy is more finely-layered in our preferred model for RAR than are the isotropic wavespeeds, suggesting complex textures. Although the halfspace in the 1-D model for RAR starts at 41-km depth (Table 2), significant depression of isotropic wavespeeds extends only to 26-km depth. (Olugboji and Park [2016] reported underplating at RAR to 22 km depth.) Deeper anisotropic layers in the model have wavespeeds characteristic of mantle rock. The particular model that we present is non-unique, but suggests some required attributes for a realistic model for the RAR RFs. Three sharp positive-polarity Ps phases with delay times < 4 sec are modelled with velocity increases at 2.8-, 12-, 18-, and 26-km depths, consistent with the volcanic layer, the oceanic Moho, the base of the underplated layer inferred by Leahy and Park [2005], and a deeper (18-26-km) layer. Complex anisotropic layering is likely a robust feature. We base this statement on the

observation that multiple RF pulses indicate multiple interfaces (of some kind) under the island, and that the “Unmodelled” stack of radial and transverse RFs tends to be smaller than the stack that would enhance an anisotropic signal. Simple strain models, such as distributed shear through a layer, are unlikely. The anisotropic “slow” symmetry axes within layers tend to strike E-W to NE-SW (Figure 5), while fast-axes strike N-S to NW-SE, suggesting common stress/strain orientations during rock-texture development. The layers manifest slow-axis and fast-axis anisotropy somewhat randomly, however, perhaps indicating that the true anisotropy has orthorhombic, or lower, symmetry. One concern is that anisotropic layering could be fractal, with large variance at all length scales. This appears not to be the case for RAR. Nearly all the features in Figure 3 persist as the cutoff frequency is reduced to

fc=2.0 Hz, albeit in smoother form, and new features do not appear when the cutoff ©2019 American Geophysical Union. All rights reserved.

frequency is raised to fc=6.0 Hz. (See Supplementary Information and Figure 10 below for a

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similar lowpass comparison involving station POHA.) 2.3: Fine-Layering With Limited Reverberation: Station POHA

For GSN Station POHA, we selected 994 earthquakes from the 1999-2016 interval, using

only the 00 location code at 20-Hz sampling with a KS-54000 borehole sensor (Figure 6). Although usable seismic events are numerous, Hawaii’s isolation within the Pacific plate

limits our ability to detect epicentral moveout (Figure 7). Evidence for a shallow reverberation can be found at roughly 2.2- and 3.0-sec, where a large positive pulse and a small negative pulse arrive roughly at times consistent with predicted resonances of a direct Ps arrival at 0.7-sec delay time. As the synthetics in Figure 7 indicate, the positive and negative reverberations should be equal in amplitude, which argues that an additional direct Ps phase contributes to the signal at 2.0-2.2-sec delay time. Comparison of real-data and synthetic RFs suggests that reverberations of this inferred 2-sec direct Ps phase are not observed in data RFs. The harmonic expansion of the POHA RFs displays large 2-lobed cos() and sin() terms

in which the Anisotropy/Dip combination of radial and transverse RFs is far larger than the Unmodelled combination (Figure 8). The back-azimuth variation of the POHA RFs are plotted in narrow bins of back azimuth in Figure S6, estimated for cutoff frequencies fC =1.0 Hz, 2.0 Hz, 4.0 Hz and 6.0 Hz. For the harmonic-regression RF estimates (Figure 8), the Anisotropy/Dip combination for the 4-lobed cos(2) term also predominates the Unmodelled combination, but this is not true for the sin(2) case. Therefore we forward-modelled the POHA RFs primarily against 2-lobed variation with back azimuth (Figure 9). The constant term of the synthetic RFs includes all reverberations induced by model interfaces. It therefore exhibits more negative-amplitude features than does the data-derived constant-term trace at POHA (Figure 8). The model specifies mantle-like wavespeeds below 32-km depth, in agreement with Olugboji and Park [2016], overlain by a thick (21-32-km) layer of anisotropic rock of intermediate wavespeed. Because we did not attempt to fit the 4-lobed cos(2) and sin(2) RFs in detail, we could restrict the model to slow-axis anisotropy. The slow-axis restriction should not be taken as strong independent evidence for this geometry, however. The strike of the symmetry axis lies roughly N-S in the shallowest layers, and NE-SW at depth. The strongest Vp anisotropy (10% or more) is inferred for the shallowest layers (Figure 9), with orientation that resembles the strike reported by Olugboji and Park [2016] for

a single anisotropic layer beneath POHA. Weak anisotropy is prescribed in the 5-15-km

©2019 American Geophysical Union. All rights reserved.

depth interval, corresponding to the oceanic crust. RFs for both RAR and POHA indicate the

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presence of strong anisotropy in layers beneath the oceanic Moho. The sharpness of pulses that comprise the fC=4.0Hz synthetic RFs is not matched by the broader Ps pulses in the fC=4.0Hz RFs for POHA data, suggesting that the layer interfaces are gradational. When

lowpassed to fC=2.0 Hz (Figure 10), the RFs for POHA synthetics and data bear greater resemblance, buttressing this suggestion. 2.4: Data-Analysis Discussion We estimate high-frequency harmonic regression of receiver functions above from only

two stations, but these confirm and extend the previous suggestion from Olugboji and Park [2016] that anisotropy is common in the underplated layers beneath hotspot islands. The model we infer for POHA possesses a thick underplated layer below the nominal oceanic Moho at the 17-19-km transition of Vp from 6.6 km/s to 7.2 km/s (Table 2), but RAR is less simple. There is a clear indication that the layering of seismic anisotropy is complex, though not fractal, because the number of likely layer interfaces in the RFs do not increase for MTC RF cutoff frequencies fc>2 Hz. The 1-D seismic models that we tabulate in Tables 2 and 4 for RAR and POHA, respectively, are illustrative rather than definitive. Our assumption of anisotropy has a symmetry axis may be too simple. It also seems likely that shallow interfaces directly beneath the stations are sufficiently planar to generate a coherent direct Ps arrival, but not planar enough to focus reverberations between buried interfaces and the seafloor. Because our reflectivity synthetics contain these reverberations, at both stations the constant-back-azimuth RFs are impossible to fit well. For example, we interpret the direct Ps phases at 2.3-2.4-s time delay in RAR and POHA data RFs (Figures 3 and 8) as converted

from the oceanic Moho. In both POHA and RAR synthetic RFs (Figures 5 and 9) these Moho-Ps pulses for the “constant” back-azimuth stack are combined with reverberations from shallower interfaces and do not match the data RFs well. The forward modelling of the Moho structure is guided instead by the harmonic RFs stacks that vary with sine and cosine of the back azimuth. Synthetic seismograms indicate that reverberations have much smaller expression in these harmonic stacks. Some evidence for shallow anisotropy surrounding mid-ocean volcanic islands can be

found in earlier studies. Levin and Park [1998b] reported Love-to-Rayleigh converted

surface waves observed at GSN station KIP at Kipapa, Hawaii, likely caused by an abrupt lateral gradient somewhere within the lithosphere, located 3°-5° from the station. This inferred change in anisotropic fabric would lie near the edge of the Hawaii swell. A similar change in anisotropic fabric at the boundary of the Kerguelen Plateau was inferred from ©2019 American Geophysical Union. All rights reserved.

surface-wave polarization anomalies by Pettersen and Maupin [2002]. At the opposite end of

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the seismic spectrum, Raitt et al. [1971] reported 8% Vp anisotropy from head waves in

marine-refraction surveys on the Hawaiian swell (called the Hawaii “Arch” at that time) whose fast orientation deviated 20°-30° from the nearby Molokai Fracture Zone. This anisotropy was both the largest and deviated the most from the fossil spreading direction, out of 7 locations surveyed by Raitt et al. [1971] across the Pacific Ocean. It therefore is consistent with a sub-Moho layer that developed its anisotropy as a byproduct of hotspot activity. The mid-20th-century Hawaii refraction experiments lay within the underplated area reported by Leahy et al. [2010], so refraction-survey headwaves would have been influenced by anisotropy within the underplated layer. Recent active-source surveys reported no underplated layer [Ohira et al., 2018] across a transect of the North Arch, northward of where the Raitt et al. [1971] surveys reported sub-Moho anisotropy and the OBS sites where Leahy et al [2010] reported an underplated layer. This suggests that the underplated layer does not extend across the entire Hawaii swell. 3: Serpentinization and Plate Tectonics A underplated layer has been detected under many, if not most, hotspot islands. If the

underplated layer developed as a simple pluton, it should have minimal anisotropy, because its mineral orientations will be random. A layered pluton or a stack of injected sills would likely be dominated by vertical-axis anisotropy, which does not induce P-SH converted waves that diagnose anisotropy in receiver-function studies. How could sub-Moho serpentinization lead to anisotropy in the underplated layer? Serpentine is rare in mantle xenoliths, but Boudier et al. [2010] describe examples that suggest its formation along fracture networks within an ultramafic protolith. Such rock textures would lead to CPO anisotropy (crystallographic-preferred orientation) independent of any serpentine LPO

(lattice-preferred orientation), owing to the distinct elastic properties of serpentine and its protolith minerals [Morales et al., 2013]. The hypothetical scenario that we present in this section agrees with Olugboji and Park [2016] that anisotropy orientations in the underplated

layer could reflect stresses that governed sub-Moho fracturing during hotspot formation. The oriented fractures facilitate seawater infiltration in this paper, rather than magma ascent. The acronym for our conceptual model is MSUP, for “metasomatic underplating.” In this

scenario a thermal plume head at the base of the lithosphere need only supply fresh partial melt upwards in a small area of active hotspot volcanism, rather than across the full width of the hotspot swell. The hotspot swell, in turn, would be maintained partially as a metasomatic ©2019 American Geophysical Union. All rights reserved.

“inflammation” of the mantle, like a pimple, rather than (primarily) as a dynamically-

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supported thermal plume. If the plume swell is entirely metasomatic, we lose one of the five criteria that Courtillot et al. [2003] identify as necessary conditions for a deep-seated mantle plume, in the sense of Morgan [1972]. Three of the remaining four criteria remain satisfied for Hawaii: a linear track of past volcanic centers, low Vs at 500-km depth, and high 3He/4He

ratio. The fifth criterion is a burst of plume-head magmatism, but this feature would have subducted under Asia by now. Later in this paper we outline reasons to believe that the plume-swell support remains largely thermal and/or dynamic. A metasomatic origin for the underplated layer would reframe the plume hypothesis for Hawaii, but not bury it. The serpentinization of shallow mantle has been invoked many times to explain isostatic

imbalances, tectonic uplifts and seismic-velocity deficits, though not always successfully. Hess [1955] suggested that serpentinization could explain “suboceanic topography,” but with water sourced from mantle degassing. An elaboration of this proposal occupies much of Hess [1962] as well, the famous “essay in geopoetry” that nucleated plate tectonics. Hess’s mantledegassing proposal was discarded. Serpentinization from seawater cooling at the mid-ocean

ridge is unmentioned in the seafloor-spreading review by MacDonald [1982], but was revived by Cannat [1993] for slow-spreading ridges. Moho temperatures at fast-spreading mid-ocean

ridges are typically too high for serpentine formation [Maclennan et al., 2005; Faak and Gillis, 2016]. Widespread serpentinite is produced beneath the seafloor at slow-spreading ridges [Charlou et al., 1998, Dilek et al., 1998] and has been inferred to underlie some

oceanic fracture zones [Detrick et al., 1993; Calvert and Potts, 1985]. Keith [2001] reasoned from several sources to assert that aseismic ridges, such as the NinetyEast Ridge in the Indian Ocean, are buoyed by serpentinized mantle, but reasoned further that active-spreading-ridge low-velocity zones were hydrated mantle. This extension put Keith [2001] in conflict with the plate-tectonics paradigm, which predicts partial-melt zones to supply the ridge [e.g., Cochran and Buck, 2001]. Despite these failures, there are many accepted roles for serpentinization in the plate-

tectonics paradigm. Serpentine is a key mineral for storing and transporting fluids that ascend from the top surfaces of subducting slabs [Hacker et al., 2003]. Faccenda et al. [2008] modelled cracks that develop in slabs as they bend downward in subduction zones, allowing water to hydrate the mantle [Rupke et al., 2004; Emry and Wiens, 2015] and to generate anisotropy with a trench-parallel fast polarization, as observed in most shear-wave birefringence surveys [Long and Silver, 2008]. Peacock [2001] argued that deeply-hydrated

slabs are responsible for double seismic planes in shallow subduction zones. Bayrakci et al. ©2019 American Geophysical Union. All rights reserved.

[2016] argue that volumes of low-velocity upper mantle offshore Iberia represent

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serpentinized peridotite that was exposed to seawater that percolated down normal faults from the opening rifts of the North Atlantic Ocean, see also Wu et al. [2006] and Ros et al.

[2017]. At the largest scale of the Earth system, Korenaga [2011] argues that ongoing

hydration of Earth’s mantle from slab-transported seawater, in whatever form, is necessary to maintain mantle convection and plate tectonics from the PreCambrian to the present. Hotspot geodynamics, in fact, is one of the few major plate-tectonics processes in which serpentinization has not been proposed as a key factor. Several serpentinization reactions are typically highlighted in the literature [Bach et al.,

2006; Seyfried et al., 2007; Kelemen et al., 2011; Humphris and Klein, 2018]. However, as

pointed out by several authors [Thayer, 1966; Hostetler et al., 1966; Moody, 1976; Evans et al., 2013], the overall formation of serpentine from olivine is best described without stoichiometric coefficients as the hydration-hydrolysis reaction Mg-Fe olivine + H2O = MgFe serpentine + Mg-Fe brucite + magnetite + H2. The stability fields of serpentine and olivine do not depend strongly on iron content, at least for olivine that contains less than 10 wt-% Fe. Because the solid-solution chemistry for serpentinization is not well understood, it is customary to use pure-Mg olivine and serpentine to define the stability fields of olivine and

serpentine in T-P space. Figure 11 shows the stability fields of minerals in the pure MgO-SiO2-H2O system

between pressures of 0 and 10 kbars and temperatures between 0 and 600°C [Evans et al, 2013] for all regions below the T and P of the reaction MgSiO4 + H2O → Mg3Si2O5(OH)4 +

Mg(OH)2, serpentine will form from Mg olivine if sufficient H2O is delivered to the reaction site. At the atomic scale, brucite (Mg(OH)2), serpentine (Mg3Si2O5(OH)4) and talc (Mg3Si4O10(OH)2) are mineralogic cousins, transitioning from a mineral composed of Mg-

hydroxyl sheets (brucite), to Mg-hydroxyl sheets that alternate with single sheets of silica tetrahedra (serpentine) to one sheet of Mg-hydroxyl paired with two silica sheets (talc) [Evans et al., 2013]. If seawater acquires dissolved silica during its descent through the basalts and gabbros of the oceanic crust, at first more serpentine would likely be produced from olivine than brucite, and talc may also appear as a product. Aside from mantle hydration, such a process would also raise the silica content of the sub-Moho mantle, leading to broader impacts we discuss in a companion paper [Park and Rye, 2019]. Figure 12 alters the phase diagram shown in Figure 11 with depth relations shown on the

right-hand axes assuming ocean depths of 500 meters (Figure 12a), and 5000 meters (Figure 12b), and a basaltic crust thickness of 8 kilometers to the top of the Moho. Superimposed on ©2019 American Geophysical Union. All rights reserved.

the diagram are average geothermal gradients ranging from 10°C/km to 30°C/km. Note that

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in the open ocean the mantle geotherm is most likely closer to 10°C/km. For parts of the ocean crust that are proximal to active magmatism the geotherm is likely to be higher, with 35°C/km a rough upper bound. Most of the seafloor will lie sufficiently far from active magmatism to overlie a 10-km layer of upper mantle that will fall within the stability field of serpentine (typically lizardite or antigorite), assuming that H2O can be delivered. However,

wherever average geothermal gradients ≧30°C/km, the depth that serpentine can be made is restricted to less than 6 km. Temperatures may be too high for sub-Moho serpentinization at most mid-ocean spreading ridges, or beneath hotspot islands that are located at spreading ridges. We conclude that, with delivery of enough H2O through the oceanic crust of a hotspot swell, serpentine can be made in the upper mantle in contact with overlying basaltic crust that is neither too thick nor too young, except proximal to active magmatism. Volcano

seismicity typically indicates that main-stage hotspot magmas are delivered via a fairly narrow (~5 km) conduit from the asthenosphere [e.g., Lin and Okubo, 2016; Klein, 2016]. Narrow plume conduits will restrict the area of the Moho that is too hot for serpentinization to proceed.

4: Paradigm Shootout: Plume Plutonism Versus Metasomatism 4.1: The Case for Magmatic Underplating An igneous origin for ocean-island underplating is supported by the overall agreement of

many geophysical observables with the classic plume model. Whole-mantle tomography is increasingly capable of resolving low-velocity features that confirm the hypotheses of ascending plumes from the core-mantle boundary, matched with locations of major mid-plate volcanism [French and Romanowicz, 2015]. In the case of the largest hotspot swell, multiple recent studies have confirmed the existence of low-velocity upper-mantle beneath the Hawaiian chain, with plausible connections to a deep plume source [Laske et al., 2011; Wolfe et al., 2009; 2011; Cheng et al., 2015]. The Sp receiver-function study by Rychert et al. [2013] identifies a velocity interface at 100-150-km depth that they identify with the onset of

melting as the hot plume rises through the pressure-dependent solidus. In geodynamical modelling, the idealized plume models of Ribe and Christensen [1994; 1999] and Yamamoto and Morgan [2009] can explain many features of Hawaiian swell bathymetry and island volcanism. Later studies of plume bathymetry have confirmed the general consistency of the plume model in terms of buoyancy flux and dynamic topography [Crosby and McKenzie,

©2019 American Geophysical Union. All rights reserved.

2009; King and Adam, 2014; Wessel, 2016]. With regard to widespread crustal underplating

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of the Hawaiian swell reported by Leahy et al. [2010], Cadio et al. [2012] argue that the admittance between geoid and topography across the plume argues for a compensation depth deeper than this sub-Moho layer. Finally, Richards et al. [2013] present a petrologic model that matches the Leahy et al. [2010] argument that magmatic underplating must be sourced over a broad region below the Hawaii swell, rather than spreading laterally as a diapir or sill. A key aspect of the Richards et al. [2013] model is the generation of ultramafic primary melts via pressure-temperature conditions unique to mature oceanic lithosphere at 60-90-km depth, formulated to pond at the base of the oceanic crust, rather than to rise through it. An outstanding paradox of the plume model is that heat flow does not correlate with

bathymetric features according to simple thermal models [e.g. Crough, 1983]. The lack of heat-flow variation across hotspot swells [von Herzen et al., 1989; DeLaughter et al., 2005] is

inconsistent with thermal erosion at the base of the lithosphere, leaving the hotspot swell to be supported by buoyancy flux or compositional variations [Foulger, 2012; Stein and Stein, 2015]. Similar conclusions can be drawn from correlations between residual topography and gravity in global-ocean bathymetry [Crosby and McKenzie, 2009; Tondi et al., 2017]. Leahy et al. [2010] endorsed the compositional hypothesis in the form of their widespread underplated layer beneath the Hawaiian Swell. Neither Leahy et al. [2010] nor the more elaborate petrological scenario of Richards et al. [2013], however, estimates how much the magmatic intrusion of a 5-10 km underplated layer would enhance the heat flow of the swell, perhaps worsening the heat-flow paradox. Harris and McNutt [2007] argue from the large spatial scatter of seafloor heat-flow measurements that the plume heat-flow paradox can be attributed to cooling of the oceanic crust by hydrothermal circulation. This would concentrate lithospheric heat flow, they argue, into restricted areas that could be undersampled by seafloor surveys. We note that hydrothermal circulation as proposed by Harris and McNutt [2007], if it were to penetrate through the crust into the mantle, has the potential to trigger the serpentinization scenario for metasomatic underplating that we articulate in this paper. Harris

and McNutt [2007] did not specify locations of hydrothermal circulation; they suggested that thick sediment cover could mask its operation beneath the seafloor. Geodynamic models of plume swells typically posit a large, warm, reduced-density body

beneath a thick lithospheric plate that dampens any enhancement in heat flow; DeLaughter et al. [2005] suggest thicknesses that range between 75 and 125 km. Any viscous flow associated with underplating of the thin oceanic crust of a hotspot swell, such as proposed by ten Brink and Brocher [1987], would plausibly exert stresses that would influence ©2019 American Geophysical Union. All rights reserved.

bathymetry. A similar scenario was hypothesized to explain Venus coronae (Figure 13),

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which can manifest as topographic rings with 0.5-1.5-km peak-to-peak relief [Squyres et al. 1992; Koch and Manga, 1996; Grindrod and Hoogenboom, 2006]. Koch [1994] and Koch and Koch [1995] modelled Venus coronae as the response of a free-slip surface to the rise and spread of buoyant, viscous diapir, evolving from dome-like novae features to coronae,

that is, large circular rings with paired extensional and compressional stresses in the radial direction. Koch and Manga [1996] extended this modeling to allow the rising diapir to stall at a depth of neutral buoyancy, as a dense mafic-to-ultramafic intrusion might stall beneath the Moho, similar to the scenario proposed by Richards et al. [2013]. This theoretical extension successfully replicated the circular-ridge topography of Venus coronae when the aspect ratio of the spreading diapir reached roughly 8 [Koch and Manga, 1996]. This critical ratio corresponds to the horizontal radius of the coronae reaching roughly twice that of a diapir that was originally spherical. Reviews of corona models by Grindrod and Hoogenboom [2006] and Jurdy and Stoddard [2007], and the addition of thermal effects by Gerya [2014],

indicate that the spreading-diapir model remains a physically-reasonable explanation for the topography of Venus coronae and other plume-related features. Ring-topography is not observed in the bathymetry surrounding mid-ocean hotspots.

However, certain features of Venus coronae appear to be replicated by the Yellowstone hotspot. Mantle-derived basalts have intruded into thick North American continental crust to form an age-progressive chain of silicic volcanoes since 16 Ma [Smith and Braile, 1994]. Shervais and Vetter [2009] and Szymanowski et al. [2015] argue that the Yellowstone hotspot has left a trail of basaltic flows, mostly tholeiitic, in the Snake River Plain. The geophysical imprint of the plume within the crust includes the collapsed Snake River Plain (SRP) along the volcanic track surrounded by a parabolic bow-shock of earthquake activity bordering the broad Yellowstone caldera, extending 100s of km through the surrounding Rocky Mountains [Anders et al., 1989]. Seismic tomography images a low-Vp region that Farrell et al. [2014] interpret as a broad magma reservoir roughly 10 km below a 50-km segment of the SRP adjacent to Yellowstone caldera. Although Yellowstone’s silicic volcanic petrology differs from an oceanic hotspot and (probably) a Venus corona, the depth and lateral extent of the SRP magma reservoir are similar to that expected for Venus and for magmatic underplating beneath hotspot islands smaller than Hawaii. Because Venus is a oneplate planet, its plume heads do not form tracks, so corona surface deformation is circular or elliptical, rather than parabolic. The sunken topography of the SRP magma reservoir (within the region uplifted by the broader Yellowstone plume) and the active deformation within the ©2019 American Geophysical Union. All rights reserved.

upper crust at the edges of the reservoir would find analogies in the physical model of Koch

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[1994] and Koch and Manga [1996] for Venus coronae. Rising-diapir models for Venus pose quantitative difficulties for magmatic/diapiric

underplating of hotspot swells on Earth. Assuming a 7-km thickness for an underplated layer beneath the oceanic Moho, and scaling the calculations of Koch and Manga [1996], we suggest that raised circular rims should be seen in seafloor bathymetry for any magmatic or viscous intrusion that extends >56 km in diameter. Nearly all plume swells are much wider than this value [King and Adams, 2014], but no bathymetric features that resemble Venus coronae have been reported. We contend that this “dog that did not bark” [Doyle, 1894] rules out underplating via a single diapir as a universal underplating mechanism. Similarly, we can rule out the underplating of a larger swell via a composite patchwork of rising diapirs, such as Richards et al. [2013] argue would have “ponded” below the Moho after ascending via porous or vein-focussed flow from a sub-lithosphere partial-melt region. Fully-extended

diapirs of ~56-km diameter or more would have left a mesh of overlapping circular topographic features, as sometimes seen on Venus [Jurdy and Stoddard, 2007]. Underplating via smaller diapirs, more numerous but with smaller aspect ratio, would be predicted to produce domal features that are also not observed. We conclude that “diapiric” underplating is not a viable model for Earth hotspot swells. The magmatic underplating model of Richards et al. [2013] could avoid forming

substantial spreading diapirs if the ascent of its ultramafic magma from the plume head were

sufficiently diffuse. This seems unlikely, given that melt-ascent of extrusive volcanism is observed to concentrate on scales matching the spacing between hotspot islands [Hieronymus and Bercovici, 2001], or between volcanic centers within islands [Moore and Clague, 1992; Lipman and Calvert, 2013]. Diffuse magmatic underplating of a large swell, or an elevated aseismic ridge, would involve more magma than needed to form its islands and seamounts, so it would not be a secondary contribution. Diffuse magmatic infiltration that extends laterally for 100s of km across the Hawaiian hotspot swell would raise the temperature of the lithosphere towards its solidus, in conflict with seismic tomography of the swell region [e.g., Laske et al. 2011]. In addition, it seems difficult to reconcile broad and diffuse magma ascent beneath the Hawaiian swell with the localized eruption of the North Arch basalts on the swell, as described by Yamamoto and Morgan [2009]. 4.2: The Metasomatic Underplating (MSUP) Hypothesis We motivate the metasomatic underplating hypothesis by analogy to active-tectonic

activity on the seafloor where large volumes of peridotite mantle are exposed to seawater. ©2019 American Geophysical Union. All rights reserved.

Whether exposed at the crests of ultra-slow ridges [Dick et al., 2003] or within isolated

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hydrothermal fields off the slow-spreading mid-Atlantic ridge [Schroeder et al., 2002; Blackman et al., 2002; Dias et al., 2010], these areas experience intense serpentinization and localized uplift associated with detachment faulting. The conversion of olivine to serpentine in the exposed peridotites involves a ~30% solid-phase expansion, and Andreani et al [2007] estimates that 100% serpentinization of a typical peridotite increases its volume by ~27%. Schroeder et al. [2002] suggest that complete serpentinization of underlying peridotite to a depth of 3 km could have raised the Atlantis Massif, at 30°N off the mid-Atlantic Ridge, by ~1.2 km elevation relative to surrounding seafloor. (Serpentinization may only be one of

several hydration reactions to consider, e.g., Nozaka et al. [2008].) Because the Massif is no older than the young plate it sits upon, Blackman et al. [2002] estimate it to have experienced an uplift rate as large as 1.5 mm/yr. On the Atlantic Massif lies the remarkable Lost City hydrothermal vent field [Kelley et al., 2005], in which highly-reduced dissolved reactants from serpentine-formation feed methane-cycling archaea microbes at the base of an exotic ecosystem with lower biomass than typical black-smoker seafloor vents. Samples of hydrated peridotite from the Lost-City seafloor exhibit serpentinization along dense fracture networks [Schroeder et al., 2002], sometimes with olivine cores surrounded by serpentine veins in a mesh texture [Roumejon and Cannat, 2014]. Schroeder et al. [2002] suggest that the ~30% volume expansion associated with

serpentinite formation, as well as carbonate precipitation, acts to block pathways for seawater

to percolate downward. However, active-source seismic surveys of the Lost City region and selected slow-spreading ridges suggest that serpentinization penetrates several kilometers beneath the seafloor [Minshull et al., 2006]. Although Blackman et al. [2002] identified

major detachment faults around the Atlantis Massif as pathways for seawater, the serpentinization of 3-D volumes suggests a small-scale network of cracks and veins [Lowell and Rona, 2002]. DeMartin et al. [2004] describe a thermal cracking model to explain how a high-permeability network within peridotite could develop to 4-6 km depth as it cools near a spreading center. The metamorphic expansion of serpentine, relative to olivine, is also a candidate to induce such crack networks, or to maintain and extend microcrack networks induced by thermal contraction [Boudier et al., 2010; Plümper et al., 2012; Roumejon and Cannat, 2014]. Klein et al. [2015] replicated reaction-driven fracturing under laboratory conditions for the serpentinization of harzburgite. Metamorphic expansion probably contributes to forming serpentine mesh networks observed within rare xenoliths of hydrated mantle rock [Boudier et al., 2010; Smith, 2010; Manuella, 2011]. The textures of these ©2019 American Geophysical Union. All rights reserved.

xenoliths suggest that elastic anisotropy within larger volumes of serpentinized mantle has

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complex sources, with contributions from both fracture and metamorphic mesh networks and the original CPO of relict olivine minerals [Morales et al., 2013]. Aside from serpentinization at mid-ocean ridges, an additional way for water to enter an

oceanic plate is via steep normal faults that rupture as plates bend downward into a subduction zone [Rupke et al., 2004; Emry and Wiens, 2015; Cai et al., 2018]. Korenaga

[2017] expresses skepticism that seawater can channel through such earthquake faults, estimating that the deviatoric stresses supported by cracks within a permeable fault zone cannot cancel the confining pressure on its fluids sufficiently to allow further infiltration from above. Korenaga [2017] allows that thermal cracking of cooling crust would generate additional porosity and could allow fluid descent to the mantle, but judges that this effect would be limited in the forearc bulge of a subduction zone. At sites of hotspot volcanism, however, we suggest that thermal perturbations are sufficient to generate more-numerous porosity-enhancing cracks proximal to magmatic conduits than within forearc fault zones. This suggestion remains to be tested in a specific hotspot setting, even though it is now commonplace for geoscientists to assume that water can descend from Earth’s surface 10s of km into the deep crust and mantle [Rupke et al, 2004; Faccenda et al., 2008; Emry and Wiens, 2015; Bayrakci et al, 2016]. Park and Rye [2019] review geochemical evidence from hotspot volcanics, particularly late-stage eruptives, that is consistent with seawater penetration into the subsurface of the Hawaii island chain. Considering the phase relations for the antigorite polymorph of serpentine [Ulmer and

Trommsdorff, 1995], peridotite serpentinization should occur whenever a significant volume of seawater can descend through the oceanic crust through faults or cracks and encounter temperatures T<600°C at the Moho. Assuming that hydrothermal circulation at the midocean ridge ceases as new oceanic plate drifts away, the next opportunity for water to enter oceanic mantle would be where hotspot magma pierces the plate from below and erupts at the seafloor. The ascent of magma will induce stresses and small fractures in the crust, surrounding the magma conduit, perhaps re-opening the thermal-crack network that Korenaga [2017] argues to be a better choice for seawater infiltration than earthquake faults. Olive and Crone [2018] develop a thermal-cracking model for the mid-ocean ridge that cools

new oceanic crust, allowing seawater to serpentinize the mantle. If basaltic magma ponds and persists below the Moho, the hotspot mantle will be too hot to produce antigorite, so metasomatic and magmatic underplating cannot coexist.

©2019 American Geophysical Union. All rights reserved.

In the metasomatic-underplating model, once hydrothermal flow is established, the ~30%

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volume expansion associated with olivine-serpentine hydration will fracture a growing volume of peridotite mantle (Figure 14). Because proportionately more olivine is found within mantle rock than crustal rocks, the expansion of the former will exert extensional stresses within the latter. This difference in volume expansion could be exacerbated at the

oceanic Moho itself, where a layer of dunite cumulates below a gabbroic lower crust [Boudier et al., 1996] might maximize the serpentinization potential locally. Further crustpenetrating cracks form, perhaps re-activating cracks from the former mid-ocean ridge hydrothermal cooling system, porosity and permeability increase, and seawater descends to a

new patch of the oceanic Moho. The lateral growth of serpentinization is governed by the spread of thoroughgoing

fractures in the oceanic crust, adding isostatic buoyancy to the dynamic buoyancy of the plume beneath. If this crack-network growth were fully self-sustaining [e.g., Klein et al., 2015], the underplating of the Hawaii swell would consume the entire Pacific Plate over time. This is not observed, suggesting that the extensional stresses of the plume uplift are an important contribution to the expansion of metasomatism below the Moho. Metasomatic underplating is also possible beneath seafloor magmatic sources not associated with plumes,

e.g., induced by plate flexure [Hieronymus and Bercovici, 2000] or petit-spot volcanism [Hirano et al., 2016]. These weaker magmatic centers should lead to smaller patches of serpentinization than those associated with deep-seated plumes and their buoyant swells. The metasomatic underplating (MSUP) model explains several observations regarding

hotspot islands that are otherwise puzzling. Serpentinized xenoliths and seafloor peridotites display textures that promise complex anisotropic behavior in the crust and in the underplated layer beneath hotspot islands, as inferred above for GSN stations POHA and RAR. The orientation of crack networks could align partially with the ambient stress field of the moving plate, in agreement with the rough alignment of anisotropic symmetry axis and plate-motion direction reported by Olugboji and Park [2016], though it is likely that a gradual expansion of crack networks via volume-expansion stresses would also be contingent on local structures. The “inflammation” of the sub-Moho mantle surrounding a hotspot island does not involve any large-scale viscous flow that would stress the oceanic crust into ridges, as predicted by models for Venus coronae, and not observed on Earth’s seafloor. Although some heat is generated by serpentinization reactions [Schroeder et al., 2002], it would be far less heat than associated with magmatic underplating. Moreover, the tight linkage of metasomatic underplating with crust-penetrating fluid flow would likely dampen observations of seafloor ©2019 American Geophysical Union. All rights reserved.

heat flow, and help to resolve the longstanding plume heat-flow paradox [Harris and McNutt,

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2007].

4.3: Earthquake Stresses Beneath Hawaii

Different models for crustal underplating make different predictions about the stresses

beneath hotspot islands. These can be compared with observations of lithospheric seismicity, at least for the well-studied island of Hawaii. Numerous crustal earthquakes are associated

primarily with volcanic processes [Lin and Okubo, 2016], but Klein [2016] focused on earthquake mechanisms in the lithosphere, beneath the volcanic edifice in the depth range 960 km. Owing to the restricted aperture of the land-based seismic network on Hawaii Island, only axes of greatest pressure were estimated for earthquakes during 1970-2007, with Richter magnitude ≥ 2.5. Klein [2016] found marked differences in focal parameters for seismic

events above and below a horizontal plane at 21-km depth that displayed almost no seismicity. Below 21 km depth, most earthquakes displayed P-axes that aligned with radial spokes emanating from a “stress center” at the center of the island, near the Mauna Loa summit and a cluster of long-period earthquakes at 40-60-km depth that Klein [2016] identifies with the principal magma conduit for Mauna Loa. Above 21-km depth, the P-axes of lithospheric quakes are more scattered, but tend to align their P-axes with circles about this stress center, perpendicular to the radial-spoke alignment of the deeper quakes. Klein [2016] notes that Hawaii’s earthquake mechanisms are inconsistent with stresses

expected if plate flexure supports the island, e.g., the lower half of the lithosphere should be in tension. The T-axes of deeper earthquakes should be aligned with radial spokes from the

center of the load (Figure 15a), and the upper half of the lithosphere in compression, with

radially-oriented P-axes. Faced with the opposite pattern of earthquake stress release, Klein [2016] argues for a broken-plate model, with plate failure located at the stress center defined by the P-axes of the deeper earthquakes, near the inferred plume conduit, rather than along a linear feature such as the hotspot chain [Watts and Cochran, 1974] or the Molokai Fracture Zone [Wessel, 1993]. In the scenario envisioned by Klein [2016], the Pacific Plate bends downward (concave down) from roughly the coastline of the island of Hawaii toward a weak central hole (Figure 15b). The plate flexure proposed by Klein [2016] poses obstacles for the magmatic

underplating model. Compression in the lower lithosphere would impede the diffuse porousflow or fracture-guided melt migration processes described by Richards et al. [2013] and Olugboji and Park [2016]. These magma-migration scenarios would be served better if unbroken lithosphere supported the volcanic island, leaving the lower lithosphere in tension ©2019 American Geophysical Union. All rights reserved.

and susceptible to infiltration. Alternatively, a broken lithosphere might allow free ascent of

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magma through its weak spot at the plume conduit, so that magmatic underplating could spread laterally beneath the Moho. This scenario is discouraged by the lack of concentric seafloor ridges at the edge of the underplated flow, as predicted by the Venus-coronae model of Koch and Manga [1996]. One weakness of the Klein [2016] model for lithospheric flexure is that lithospheric

earthquake hypocenters beneath Hawaii do not trend deeper near the mid-island “stress

center” to follow the hypothesized plate curvature toward its point of failure; see Figure 9 of Klein [2016]. We propose an alternate interpretation of the earthquake stress-release pattern (Figure 15c). From POHA RFs, our modelled Vp (Table 4) rises from 6.6 km/sec to 7.5 km/sec between 17 and 21 km depth, consistent with a transition from gabbroic crust to partially-serpentinized mantle (~35% according to Hess [1962]) just above the neutral-stress depth identified by Klein [2016]. If we assume metasomatic underplating to currently be active in the lithosphere below the neutral plane at 21-km depth, then volume expansion associated with peridotite hydration would induce compressional stresses at broad scale, in response to the confinement of the surrounding plate, even while creating and maintaining an evolving grain-scale network of permeable cracks. Earthquake fracturing under uniform volume expansion would not necessarily exhibit the

radial orientation of the P-axes reported by Klein [2016] for earthquakes below 21-km depth, but the overlying volcanic island should orient the stresses. Hawaii island is the upper portion of a broader semi-conical edifice. Its collection of volcanoes applies a maximal vertical load within its high-altitude interior, tapering off to zero along the underwater flanks

of the edifice. The “soft center” of the Pacific plate that Klein [2016] proposes to accommodate downward bending under Hawaii could be replicated by the ductility of the serpentinized underplated layer, if it is being pushed radially away from the locus of maximal

volcanic load. Many of the earthquakes with P-axes oriented radially from the volcanic center lay deeper than the inferred serpentinized layer, including two with M>6. These deeper quakes possibly reflect either incipient serpentinization below the deepest interface of our RF-derived model, or else stresses transferred from ductile flow in the underplated layer into the brittle lithosphere immediately below. Lithospheric earthquakes shallower than 21-km reside largely in the oceanic crust,

according to our POHA velocity profile. The MSUP scenario posits extensional stresses in the crust induced by volume expansion just below in the sub-Moho mantle. The reorientation

of Hawaii earthquake stresses is consistent with this scenario. In the crust the extensional ©2019 American Geophysical Union. All rights reserved.

stress directions will be determined by the fracture network that allows seawater to descend

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to the Moho and beyond. The orientation of fractures in this hydrothermal network need not follow a simple pattern, because induced fractures will be governed partly by pre-existing planes of weakness in the plate.

In agreement with this prediction, Klein [2016] reports

more scatter in the P-axes for earthquakes in the 9-21-km depth range. Metasomatic underplating can explain the pattern of earthquake stresses with depth

[Klein, 2016] without invoking a broken plate. The neutral-stress depth in Hawaiian lithospheric seismicity lies close to the top of the underplated layer that we infer for POHA in Section 2. The crossover in P-axes near 21-km depth matches the transition from

metasomatic volume increases in the shallow, and induced extensional fracturing in the crust above. Considering that a truly “broken” Pacific Plate might have fragmented along a weak Hawaii-Emperor seamount chain many times during the past 80 Myr, the persistence of an intact Pacific Plate should count in favor of the MSUP hypothesis. 5: Predictions of the MSUP Hypothesis 5.1: Hydrothermal Circulation Is there evidence for hydrothermal circulation associated with metasomatic underplating?

Low-temperature serpentinization reactions typically generate H2 as a product, leading to fluids of reduced chemistry that can support primitive life. The Lost City Hydrothermal Field [Kelley et al., 2005] supports microbes that metabolize H2 and CH4 to support a vent macrofaunal ecosystem that lacks the sulfide-metabolizing organisms that typically inhabit mid-ocean-ridge hot-vent environments. The newest center of Hawaiian magmatism is Loihi volcano, which lies underwater SW of Mauna Loa, on seafloor that lies somewhat ahead of the hot-spot swell. At 5-km ocean depth on the south slope of Loihi lies an unusual hydrothermal system with diffuse vented water only 0.2°C warmer than ambient [Edwards et al., 2011]. This vented fluid has reducing chemistry, marked by ferrous-iron Fe2+, that is metabolized by iron-oxidizing Zetaproteobacteria to ferric-iron Fe3+ and deposited as ironoxyhydroxide layers where the vent fluids meet the O2-rich ambient seawater. Loihi is only one locale for Zetaproteobacteria; similar diffuse-vent ecosystems have been discovered globally, proximal to hot sulfide-rich vent systems [Emerson and Moyer, 2010], near hotspots

[Johannessen et al., 2017], mid-ocean ridges [Scott et al., 2015] and island-arc volcanoes [Bortoluzzi et al., 2017; Ristova et al., 2017]. Bach [2016] estimates that Fe-oxidizing bacteria represent roughly 10% of the sedimentary deep biosphere.

©2019 American Geophysical Union. All rights reserved.

The reported locations of Fe2+-rich diffuse venting at the seafloor fit the pattern of

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serpentinization that we hypothesize in this paper. We know of no evidence, however, that these diffuse-venting locations connect unambiguously to serpentinization at the oceanic

Moho and deeper. For example, dissolved ferrous iron reflects the “reducing chemistry” of serpentinization reactions [Seyfried et al., 2007], but it is not a familiar product of them. Evans [2010] argues that antigorite formation at >300°C (consistent with sub-Moho hydration of mantle peridotite) would not produce the dissolved H2 that characterizes the vent fluids from Lost City [Kelley et al., 2005], produced by low-temperature serpentinization.

Park and Rye [2019] expand on this chemistry, arguing that seawater that descends through 7-10 km of tholeiitic gabbros to reach the oceanic Moho will acquire dissolved silica that will shift serpentinization away from magnetite (Fe3O4) and H2 as reaction products. Therefore, the lack of H2 at Zetaproteobacteria vents is not a dealbreaker for the metasomaticunderplating hypothesis. 5.2: Ocean-Island Uplift Following Hess [1955], a key prediction of metasomatic underplating is hotspot uplift,

caused by serpentinizing dense peridotite beneath the Moho with mineral reactions that involve ~30% volume expansion. At mid-ocean hotspots, such as Hawaii, MSUP uplift would combine with the buoyancy of the plume, requiring a nuanced analysis to isolate the effect. For ridge-centered hotspots, a delay in metasomatic underplating may occur after the hotspot island forms, because the sub-Moho peridotite at the ridge itself is too hot for serpentinization reactions to proceed. Underplating is not observed beneath ridge-centered hotspot activity in Iceland [Staples et al., 1997] and Ascension Island [Evangelidis et al., 2004] from active-source seismic studies, and Tristan de Cunha [Geissler et al., 2017] in a receiver-function study. Geissler et al [2017] argued for 18-km crust beneath Nightingale Island, adjacent to Tristan de Cunha, from a temporary seismic station, and Ryberg et al. [2017] argued for a 13-km layer of crustal Vs localized beneath both islands from ambientnoise recorded by 24-station OBS deployed around the hotspot, so this region deserves further investigation. Ramalho et al. [2017] report that the hotspot island Santa Maria in the Azores has experienced a curious history of vertical tectonics, forming at 6 Ma on young lithosphere near the ultraslow Terceira Ridge (Figure 16). This ridge forms a segment of the undersea plate boundary between the Eurasian and African plates, 480 km from the midAtlantic Ridge. As reconstructed from uplifted Pliocene shorelines and geochronology of surface rocks, Ramalho et al. [2017] infer that Santa Maria first subsided as expected from the cooling of newly-formed lithosphere, but has experienced uplift since 3.5 Ma. Ramalho ©2019 American Geophysical Union. All rights reserved.

et al. [2017] argue the magmatic underplating since 3.5 Ma has caused the uplift. Because the

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Azores hotspot intersects the young, and likely thin, lithosphere of a divergent plate boundary, it is notable that the area of uplifted bathymetry surrounding the Azores lacks the outer rims that [Koch and Manga, 1996] argue to be characteristic of shallow underplating by a rising diapir (Figure 16). Spieker et al. [2018] report Ps receiver functions that they argue are consistent with an

underplated layer beneath Santa Maria, though fewer than 25 earthquakes comprise their dataset for the island. Their RF frequency cutoff fC=1.0 Hz is too low to resolve thin layers, but make a case for underplating beneath station ROSA on Sao Jorge island, which also lies

off-ridge. Spieker et al. [2018] report no evidence for underplating beneath Terceira island, which lies on its eponymous ridge. Their analysis of 43 events recorded at station CMLA on ridge-centered Sao Miguel island does not distinguish separate direct-Ps conversions from multiple interfaces near the Moho, but Spieker et al. [2018] argue for a thin (<10 km) underplated later from later pulses that they interpret as Psms and Ppms reverberations. Ramalho et al. [2017] do not explain why magmatic underplating at Santa Maria would

be geographically restricted beneath the island, or why the emplacement of additional magma

was not associated with more subaerial volcanism. The metasomatic underplating model explains both features (Figure 16): from 6 to 3.5 Ma the sub-Moho mantle was too hot to serpentinize, but after some cooling and subsidence of the young plate as it drifted away from the Terceira Ridge, seawater penetrated its base and formed an underplated layer of hydrated peridotite. In the MSUP scenario, uplift would be localized to the area surrounding recent volcanism, where thermal cracking allows seawater access though the oceanic crust. Although evidence for igneous-intrusion in some locations been more direct than for

Santa Maria [e.g., Ramalho et al 2010], metasomatic underplating may have made a significant contribution to the eruptive and uplift histories of several volcanic edifices in this portion of the Atlantic Ocean: the Cape Verde Islands [Madeira et al., 2010], the Canary Islands [Anguita and Hernan, 2000] and Madeira Island [Ramalho et al., 2015]. Dickinson [1998] reports numerous “makatea” islands among the Cook-Austral seamount chain, reeffringed islands that uplifted by 30-100 m or more after cessation of volcanism, but hypothesizes that many makatea have been lofted by the flexural bulges of younger volcanic islands nearby. The oldest island of the Samoan hotspot, Savai’i, with shield-stage volcanism between 5.3 and 2 Ma [Workman et al., 2004], is also the highest. Konter and Jackson [2012] argue that the older Samoan hotspot islands have climbed a flexural bulge as they migrated oblique to the northern terminus of the Tonga Trench. However, the 3-D flexural model ©2019 American Geophysical Union. All rights reserved.

reported by Konter and Jackson [2012] for this complicated region does not distinguish the

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tallest Samoan Islands (Savai’I and Upolu) as having uplift rates higher than the younger, lower edifices Tutuila and Papatua. 5.3: Seamount Clustering We describe one final shootout between the magmatic and metasomatic underplating

hypotheses. The South Pacific Superswell is a broad area of uplifted seafloor [McNutt, 1998] extending over 107 km2 in area, encompassing numerous volcanic islands and seamounts that date (mostly) from mid-Cretaceous eruptions that often are not age-progressive [Hieronymus and Bercovici, 2000]. South Pacific Superswell volcanism is typically accompanied by an underplated layer [Caress et al., 1995; McNutt and Bonneville, 2000], typically assumed to be magmatic in origin [Jordahl et al., 2004; Leahy and Park, 2005], but direct evidence is scant. Hillier [2007] observed that unrelated seamounts tend to cluster in similar patches of seafloor over time, particularly within French Polynesia (within the South Pacific Superswell) and within the Darwin Rise, another superswell feature identified by McNutt [1998]. Hillier [2007] speculates that “magmatic throughput leaves the [oceanic] lithosphere more susceptible to the passage of future melts.” Magmatic underplating of superswell hotspot islands (and perhaps seamounts), by itself,

would not leave the oceanic lithosphere more susceptible to the generation of fresh magma,

because the melting point of a gabbroic pluton would likely equal that of any upwelling magma from a subsequent thermal center. By contrast, metasomatic underplating would lead to a scenario that was (remarkably) anticipated by Hess [1955] for the “Mid-Pacific Mountain Range guyots between the Marianas and Hawaii,” in other words, for the Superswell volcanic centers. Hess suggested that a serpentinized layer beneath the Moho of these features would

facilitate de-serpentinization by any later volcanism, releasing water to facilitate the volcanic output necessary to form the seafloor “mountains.” Hess [1955] proposed that a serpentinized layer forms beneath the oceanic Moho from water expelled from an upwelling mantle; in this paper we propose that serpentinization is facilitated by seawater that leaks into upper-mantle peridotite via thermal cracking from the initial stages of hotspot volcanism. If the serpentinized layer forms after the initial volcanism, its presence would facilitate partial melting from any subsequent thermal event, and therefore help to explain the seamount clustering reported by Hillier [2007]. Similarly, a serpentinized sub-Moho layer would facilitate the abundant small-scale non-hotspot seafloor volcanic features reported in French Polynesia by Hirano et al. [2016].

©2019 American Geophysical Union. All rights reserved.

Minor- and trace-element features of hotspot chemistry may help resolve whether

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Polynesian volcanism is assisted by shallow hydration of the Pacific lithosphere. Park and Rye [2019] review geochemical evidence from Hawaii that may reflect interaction between

plume melts and a metasomatized lithosphere. The chemical signature of the MSUP layer, however, will not be confined to the upper half of the oceanic lithosphere. Plate tectonics will subduct the MSUP layer and recycle it within Earth’s mantle, where it can influence deep-

seated plume sources. Park and Rye [2019] present evidence that subduction of MSUP material can help explain the “pyroxenite” signature inferred from alkali-basalt volcanism in the late-stage eruptions of large hotspots and the main stages of secondary hotspots, in the Courtillot et al. [2003] nomenclature. 6: Summary

The manifestation of mantle plumes within the plate-tectonics paradigm is conceptually

simple in the context of thermal-convection simulations, but the plume model has led to predictions that have often not been verified [DeLaughter et al., 2005; Anderson, 2011]. The model for hotspot underplating as an intrusive reservoir of plume material that supplies seafloor volcanism has intuitive appeal, but fails diverse tests of its predictions regarding seafloor bathymetry, earthquake distribution, inferred stresses, and heat flow. At the same time, seismic receiver functions suggest that the crust and underplated layer at mid-ocean hotspots have complex multilayered anisotropy that belies an emplacement mechanism via simple geodynamic flow. We propose instead a process in which hotspot volcanism represents a principal avenue

for seawater to reach the oceanic Moho and to metasomatize its peridotite mantle. Barred by high temperatures from reaching the oceanic Moho at all but the slowest-spreading oceanic ridges, underplated layers form where the mantle is cool enough to serpentinize, starting initially close to the volcanic conduit, where thermal cracking opens passageways through the oceanic crust for two-way transport of magma up, and seawater down. Perhaps facilitated by a layer of dunite cumulates at the Moho [Boudier et al., 1996], differential volume expansion

of the mantle and crust induces further cracking in the crust, inducing the lateral spread of fluid infiltration and metasomatism, reaching toward the edge of the surface extensional stresses induced by plume buoyancy flux. Section 4 compared predictions of the metasomatic model for hotspot-island underplating

against selected observations where a magmatic or diapiric model for the underplating seems problematic. There are many longstanding observations for which the traditional plume ©2019 American Geophysical Union. All rights reserved.

model for hotspot volcanism predicts behavior that is not observed, e.g., elevated heat flow

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and circular seafloor ridges, similar to the “dog that did not bark” of Doyle [1894]. The metasomatic underplating model appears to have better consistency with such observations, or non-observations. We suggest that the hypothesis deserves greater study.

Acknowledgements. This work was partly supported by Yale-NUS College, where J.

Park found a quiet office for brainstorming during Fall 2016. The authors thank several colleagues for helpful conversations during the gestation: Ruth Blake, Tolulope Olugboji, Zhen Liu, Brian McAdoo, Roy Schlische, Steven D’Hondt, Gregory Ravizza, Jay Ague, Shun-Ichiro Karato, Seth Stein and Neil Ribe. Thoughtful and detailed reviews from Ricardo

Ramalho, Tim Minshull, Wolfram Geissler, Francesco Pio Lucente, and an anonymous referee improved the manuscript significantly. We thank editor Thorstein Becker for managing the review process well. Seismic data used in this study can be obtained from IRIS Data Services (http://ds.iris.edu/ds/). E. Bibliography

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FIGURES:

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Figure 1. Locations of 302 earthquakes from 1999-2016 used to estimate receiver functions for station RAR (Rarotonga Island) of the Global Seismogtaphic Network (GSN). The backazimuth sector bounded by green lines to the west of RAR defines a dataset used for the epicentral sweep shown in Figure 2.

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Figure 2. Epicentral sweep of receiver functions to the west of GSN station RAR (Rarotonga Is.). The RFs have frequency cutoff fc=4.0 Hz. Frequency-domain RFs are binned every 5° of epicentral distance  over 10° intervals, then inverse-transformed, so that each time-domain RF trace shares some data with the traces immediately above and below it. The radial and vertical components are rotated to the LQ coordinates appropriate for the global model IASPEI91 [Kennett and Engdahl, 1991]. IASPEI91 overcompensates the true incidence angles at RAR, so that the zero-delay pulse (black line) is negative. The predicted times of the Ps, Ppms and Psms converted waves and reverberations are drawn for the three buried interfaces of the stacking model in Table 1. The left panel plots data-derived RFs; the right panel plots RFs estimated from reflectivity synthetics [Levin and Park, 1997] for the stacking model (Table 1) using the same epicentral distribution. The depths of the interfaces, 4.0 km (gold lines), 9.5 km (purple lines) and 17.0 km (green lines) are chosen to match three inferred Ps converted phases at ~0.7-sec, 1.3-sec, and 2.2-sec delay time, respectively, in the RFs estimated from RAR data. In RFs estimated from synthetic seismograms, the Ppms and Psms reverberations of the 4.0-km interface obscure the Ps converted waves from the 17.0-km interface. No such interference is evident in the data-derived RFs.

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Figure 3. Harmonic regressions over back-azimuth for receiver functions for GSN station RAR with no moveout corrections for interfaces at depth. The RFs have frequency cutoff fc=4.0 Hz. The radial and vertical components are rotated to the LQ coordinates appropriate for the global model IASPEI91 [Kennett and Engdahl, 1991]. IASPEI91 overcompensates the true incidence angles at RAR, so that the zero-delay pulse (green line at 0-sec time delay) is negative. Bootstrap uncertainties are indicated by a green margin surrounding the RF traces. The green line at 5-sec delay bounds all the likely direct Ps pulses associated with underlying crustal and underplated structure. The left panel plots RFs in which the phasing of radial and transverse RFs follows the first-order theoretical effects of anisotropy and dipping interfaces [Park and Levin, 2016b]. The right panel plots RFs in which the phasing of radial and transverse RFs incorporates random Ps scattered waves and second-order effects of anisotropy and dipping interfaces.

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Figure 4. Harmonic regressions over back-azimuth for receiver functions for GSN station RAR, with moveout corrections for interfaces at 30-m depth (left panel) and 70-km depth (right panel) according to the moving-window method of Helffrich [2006] and Park and Levin [2016a]. The RFs have frequency cutoff fc=4.0 Hz. The radial and transverse RFs are combined with the back-azimuth phase that is predicted to correspond to the first-order effects of anisotropy and dipping interfaces. Bootstrap uncertainties are indicated by a green margin surrounding the RF traces. The green line at 0-sec time delay corresponds to hypothetical Ps pulses from interfaces at 30-km and 70-km depth, respectively, in the context of the stacking model for RAR (Table 1). The green line at -4.0-sec (left) and -8.2-sec (right) delays mark the zero-delay direct P wave, which is defocused by the moveout corrections here.

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Figure 5. Harmonic regressions over back-azimuth for receiver functions for synthetic seismograms at GSN station RAR with no moveout corrections (left panel), calculated for an anisotropic 1-D model (Table 2) with specified isotropic Vp, Vs and density  (upper right panel), and symmetry axis with specified tilt (middle right panel) and strike angles (lower right panel). The RFs have frequency cutoff fc=4.0 Hz. Symbol size in the lower-right panels indicates anisotropic strength. Blue up-triangles indicate fast symmetry axes; red downtriangles indicate slow symmetry axes. The green line at 4.0-sec delay lies between direct-Ps arrivals from interfaces at 26 km and 35 km. The “constant” trace contains several negative pulses that correspond to Psms reverberations in the 1-D reflectivity synthetics. Evidence for such reverberation is scarce in the data-derived RFs, see Figures 3 and 4.

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Figure 6. Locations of 994 earthquakes from the 1999-2016 interval used to estimate receiver functions for station POHA (Pohakuloa, Hawaii Island) of the Global Seismographic Network (GSN). The back-azimuth sector bounded by green lines to the west of POHA defines a dataset used for the epicentral sweep shown in Figure 7.

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Figure 7. Epicentral sweep of receiver functions to the west of GSN station POHA (Pohakuloa, Hawaii Island). The RFs have frequency cutoff fc=4.0 Hz. Frequency-domain RFs are binned every 5° of epicentral distance  over 10° intervals, then inverse-transformed, so that each time-domain RF trace shares some data with the traces immediately above and below it. The radial and vertical components are rotated to the LQ coordinates appropriate for the global model IASPEI91 [Kennett and Engdahl, 1991]. IASPEI91 overcompensates the true incidence angles at POHA, so that the zero-delay pulse (black line) is negative. The predicted times of the Ps, Ppms and Psms converted waves and reverberations are drawn for the three buried interfaces of the stacking model in Table 3. The left panel plots data-derived RFs; the right panel plots RFs estimated from reflectivity synthetics [Levin and Park, 1997] for the stacking model using the same epicentral distribution. The depths of the interfaces, 5 km (gold lines), 17 km (purple lines) and 39 km (green line) are chosen to match three inferred Ps converted phases at ~0.7-sec, 2.2-sec, and 5.2-sec delay time, respectively, in the RFs estimated from POHA data. In RFs estimated from synthetic seismograms, the Ppms and Psms reverberations of the 5.0-km interface obscure the Ps converted waves from the 17.0-km interface. A negative pulse in the data-derived RFs could plausibly be associated with the Psms phase, but is small relative to the positive pulse at ~0.7 sec corresponding to the primary Ps conversion, suggesting that the reverberation is poorly focussed.

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Figure 8. Harmonic regressions over back-azimuth for receiver functions for GSN station POHA with no moveout corrections for interfaces at depth. The RFs have frequency cutoff fc=4.0 Hz. The radial and vertical components are rotated to the LQ coordinates appropriate for the global model IASPEI91 [Kennett and Engdahl, 1991]. IASPEI91 overcompensates the true incidence angles at POHA, so that the zero-delay pulse (green line at 0-sec time delay) is negative. Bootstrap uncertainties are indicated by a green margin surrounding the RF traces. The green line at 5-sec delay bounds all direct Ps pulses associated with shallow crustal and underplated structure. The left panel plots RFs in which the phasing of radial and transverse RFs follows the first-order theoretical effects of anisotropy and dipping interfaces [Park and Levin, 2016b]. The right panel plots RFs in which the phasing of radial and transverse RFs incorporates random Ps scattered waves and second-order effects of anisotropy and dipping interfaces.

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Figure 9. Harmonic regressions over back-azimuth for receiver functions for synthetic seismograms at GSN station POHA with no moveout corrections (left panel), calculated for an anisotropic 1-D model (Table 4) with specified isotropic Vp, Vs and density  (upper right panel), and anisotropic symmetry axis with specified tilt (middle right panel) and strike angles (lower right panel). The RFs have frequency cutoff fc=4.0 Hz. Symbol size in the lower-right panels indicates anisotropic strength. Blue up-triangles indicate fast symmetry axes; red downtriangles indicate slow symmetry axes. The green line at 3.8-sec delay marks the direct-Ps arrival from model interface at 32 km, the bottom of the underplated layer. The “constant” trace contains several negative pulses that correspond to Psms reverberations in the 1-D reflectivity synthetics. Evidence for such reverberation is scarce in the data-derived RFs for POHA, see Figure 8.

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Figure 10. Data-synthetic comparison for harmonic regressions over back-azimuth for receiver functions for GSN station POHA with no moveout corrections for interfaces at depth. The RFs have frequency cutoff fc=2.0 Hz, comparing results for data (left panel) and synthetics (right panel). The radial and vertical components are rotated to the LQ coordinates appropriate for the global model IASPEI91 [Kennett and Engdahl, 1991]. Bootstrap uncertainties are indicated by a green margin surrounding the RF traces. The green line at 5-sec delay bounds all direct Ps pulses associated with shallow crustal and underplated structure. Both panels plot RFs in which the phasing of radial and transverse RFs follows the first-order theoretical effects of anisotropy and dipping interfaces [Park and Levin, 2016b]. Better agreement is evident between data and synthetic RFs for this lower frequency cutoff. Possible explanations include gradational interfaces in the real Earth, or irregular interfaces that broaden Ps pulses when moveout corrections are applied for a flat-layered stacking model (Table 3).

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Figure 11. A possible phase diagram for the system MgO-SiO2-H2O. Wide, open bands represent uncertainty in the H2O-constant reactions. Mineral relationships in the divariant fields are shown on the MgO-SiO2 binary projected from H2O. A=Atg=antigorite, B=Brc=brucite, F=Fo=forsterite, L=Liz=lizardite, T=Tlc=talc. Figure and legend after Evans et al. [2013].

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Figure 12. Phase diagram for the MgO-SiO2-H2O system, with estimated depth from the bottom of the seafloor shown, with average geothermal gradients from 10°C/km to 30°C/km superimposed on the figure. Conditions for case A (left panel) are 500-m seawater depth and 8km crustal thickness to the oceanic Moho. Conditions for case B (right panel) are 5000-m seawater depth and 8-km crustal thickness to the oceanic Moho. Note that serpentine is expected to form in the mantle, if H2O is present, from the Moho through a ~6-km thickness of upper mantle for an average geotherm of 30°C/km and to more than 20-km thickness for a 15°C/km gradient. These relationships suggest that if water is delivered to the upper mantle via thermal cracks in the oceanic crust, serpentine is likely to form.

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Figure 13. (a) Radar image from the Magellan mission of Bahet Corona in the Fortuna region of Venus (geographical coordinates 48.4°N, 0.1°E). Roughly elliptical with major and minor axes of 230 km and 150 km, Bahet Corona is surrounded by a ring of ridges and troughs, which in places cut more radially-oriented fractures. The centers of the features also contain radial fractures as well as volcanic domes and flows. Resolution of the Magellan data is 120 meters (400 feet). Image Credit: NASA/JPL (https://photojournal.jpl.nasa.gov/catalog/PIA00461). (b) Schematic of viscous-fluid model for an idealized Venus corona, after Koch and Manga [1996], who fit their viscous-flow models to coronae with diameters 200-400 km and peak-to-peak topography of 500-1500 meters.

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Figure 14. Cartoon schematic of metasomatic underplating. In the metasomatic-underplating model, once hydrothermal flow is established, the volume expansion associated with olivineserpentine hydration will fracture a growing volume of peridotite mantle. More olivine lies within mantle rock than crustal rock, so the expansion of the former will exert extensional stresses within the latter.

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Figure 15. Conceptual models of geodynamics that determine patterns of Hawaii mantlelithosphere seismicity, as reported by Klein [2016]. The standard model in which Hawaii Island is supported by an elastic Pacific Plate (a) predicts poorly the depth variation of inferred earthquake stresses. The model suggested by Klein [2016] in which Hawaii Island is supported by edge-bending of a broken plate (b) fits the patterns of seismicity better. Metasomatic underplating (c) prescribes volume increases within the upper portion of the mantle lithosphere, generating an extensional strain that resembles the strain associated with the broken-plate model.

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Figure 16. The Azores hotspot has generated ridge-centered islands that drift off-axis over millions of years. Divergent plate boundaries marked by red lines. In the case of Santa Maria Island, geologic evidence shows that the island subsided and later uplifted after its major volcanism ceased [Ramalho et al., 2017]. The metasomatic underplating model (MSUP) can explain this history if the oceanic Moho was initially too hot for serpentinization to occur, but the thermal disturbance of hotspot volcanism was large enough to generate and maintain a network of thermal cracks that allowed seawater infiltration long enough for the mantle peridotite beneath the Moho to cool into conditions favorable for serpentinization.

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Table 1. Three-layer stacking model for receiver-function moveout corrections at GSN

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station RAR (Rarotonga Is.). Layer base (km)

Vp (km/sec)

Vs (km/sec)

 (gm/cm3)

4.0

5.5

3.07

2.6

9.5

6.65

3.54

2.8

17.0

7.52

3.85

2.9

halfspace

8.35

4.4

3.36

Table 2. Layered-anisotropy model for RAR, based on trial-and-error forward-modelling. Model is graphed in Figure 5. Because receiver-function sensitivity to Vp anisotropy is far greater than to Vs anisotropy, the latter was fixed at 60-70% of Vp anisotropy. Layer

Vp

cos(2)

Vs

cos(2)



symmetry

symmetry

base

(km/sec

anisotrop

(km/sec

anisotrop

(gm/cm3

-axis tilt

-axis

(km)

)

y

)

y

)

(from Z)

strike (from N)

1.5

4.0

-10%

2.22

-7%

2.6

55

-115

2.8

5.0

-10%

2.77

-7%

2.6

55

-115

7.0

6.3

-7%

3.5

-4%

2.6

60

60

12.0

6.35

-10%

3.55

-7%

2.8

70

90

14.5

7.05

10%

3.74

7%

2.8

70

170

16.5

7.05

10%

3.74

7%

2.9

70

155

18.0

7.42

6%

3.9

4%

2.9

70

155

22.0

7.82

-4%

4.1

-2.5%

2.9

45

45

26.0

7.82

6%

4.1

4%

3.0

70

160

30.0

8.22

-10%

4.3

-7%

3.1

80

-135

33.0

8.22

-6%

4.3

-4%

3.1

80

45

35.0

8.22

4%

4.3

2.5%

3.2

90

135

41.0

8.35

-6%

4.4

-4%

3.36

80

-120

halfspac

8.35

0%

4.4

0%

3.36

0

0

e

©2019 American Geophysical Union. All rights reserved.

Accepted Article

Table 3. Three-layer stacking model for receiver-function moveout corrections at GSN station POHA (Pohakuloa, Hawaii, USA). Layer base (km)

Vp (km/sec)

Vs (km/sec)

 (gm/cm3)

5.0

5.6

3.27

2.45

17.0

6.6

3.65

2.8

39.0

7.0

3.75

3.0

halfspace

8.35

4.4

3.36

Table 4. Layered-anisotropy model for POHA, based on trial-and-error forward-modelling. Model is graphed in Figure 9. Because receiver-function sensitivity to Vp anisotropy is far greater than to Vs anisotropy, the latter was fixed at 50-70% of Vp anisotropy. Layer

Vp

cos(2)

Vs

cos(2)



symmetry

symmetry

base

(km/sec

anisotrop

(km/sec

anisotrop

(gm/cm3

-axis tilt

-axis

(km)

)

y

)

y

)

(° from Z)

strike (° from N)

2.5

4.5

-12.5%

2.7

-7.5%

2.3

60

15

5.0

5.62

-10%

3.2

-6%

2.5

60

15

7.5

6.2

-2%

3.45

-1%

2.7

30

-65

12.8

6.6

-2%

3.65

-1%

2.8

60

115

14.2

6.6

-5%

3.65

-3%

2.8

60

-115

17.0

6.6

-10%

3.65

-6%

2.8

30

-135

19.0

6.9

-8%

3.8

-5%

2.8

45

-150

21.0

7.2

-6%

3.95

-4%

2.8

60

-150

32.0

7.5

-12%

4.1

-8%

3.0

80

195

halfspac

8.35

0%

4.4

0%

3.36

0

0

e

©2019 American Geophysical Union. All rights reserved.