Submarine geology of the Hilina slump and morpho-structural evolution of Kilauea volcano, Hawaii

Ž .Journal of Volcanology and Geothermal Research 94 1999 59–88www.elsevier.comrlocaterjvolgeores

Submarine geology of the Hilina slump and morpho-structuralevolution of Kilauea volcano, Hawaii

John R. Smith ), Alexander Malahoff, Alexander N. Shor 1

Department of Oceanography, School of Ocean and Earth Science and Technology, UniÕersity of Hawaii at Manoa, 1000 Pope Road,MSB 205, Honolulu, HI 96822-2336, USA

Received 10 May 1999

Abstract

Marine geophysical data, including SEA BEAM bathymetry, HAWAII MR1 sidescan, and seismic reflection profiles,along with recent robot submersible observations and samples, were acquired over the offshore continuation of the mobileKilauea volcano south flank. This slope comprises the three active hot spot volcanoes Mauna Loa, Kilauea, and Loihiseamount and is the locus of the Hawaiian hot spot. The south flank is the site of frequent low-intensity seismicity as well asepisodic large-magnitude earthquakes. Its sub-aerial portion creeps seaward at a rate of approximately 10 cmryear. TheHilina slump is the only large submarine landslide in the Hawaiian Archipelago thought to be active, and this study is one ofthe first to more highly resolve submarine slide features there. The slump is classified into four distinct zones from nearshoreto the island’s base. Estimates of size based on these data indicate a slumped area of 2100 km2 and a volume of10,000–12,000 km3, equivalent to about 10% of the entire island edifice. The overall picture gained from these data sets isone of mass wasting of the neovolcanic terrain as it builds upward and seaward, though reinforcement by young andpre-Hawaii seamounts adjacent to the pedestal is apparent. Extensive lava delta deposits are formed by hyaloclastites anddetritus from recent lava flows into the sea. These deposits dominate the upper submarine slope offshore of Kilauea, withpillow breccia revealed at mid-depths. Along the lower flanks, massive outcrops of volcanically derived sedimentary rockswere found underlying Kilauea, thus necessitating a rethinking of previous models of volcanic island development. Themorphologic and structural evolutionary model for Kilauea volcano and the Hilina slump proposed here attempts toincorporate this revelation. A hazard assessment for the Hilina slump is presented where it is suggested that displacement ofthe south flank to date has been restrained by a still developing northeast lateral submarine boundary. When it does fullymature, the south flank may be more subject to land slips triggered by large, long duration earthquakes and thus Kilauea mayundergo more frequent episodes of failure with increased displacements. q 1999 Elsevier Science B.V. All rights reserved.

Keywords: submarine geology; morpho-structural; Kilauea volcano

) Corresponding author. Tel.: q1-808-956-9669; fax: q1-808-956-2136; E-mail: [email protected]

1 Present address: National Science Foundation, OceanographicInstrumentation and Technical Service Program, 4201 WilsonBlvd., Room 725, Arlington, VA 22230, USA.

1. Introduction

Though much has been recently learned about themorphology and extent of giant submarine landslidesfrom oceanic volcanic islands and ridges with the

0377-0273r99r$ - see front matter q 1999 Elsevier Science B.V. All rights reserved.Ž .PII: S0377-0273 99 00098-0

( )J.R. Smith et al.rJournal of Volcanology and Geothermal Research 94 1999 59–8860

Žadvent of sidescanning sonars Lenat et al., 1989;´Moore et al., 1989, 1994a,b; Holcomb and Searle,1991; Watts and Masson, 1995; Hampton et al.,

.1996 , little is known regarding their timing, trigger-ing, and mechanics of motion. Not only are mostHawaiian slide deposits located at abyssal depths andcomposed mainly of acoustically opaque lava blocksfrom the island mass, the pelagic sedimentation ratesare quite low with a relatively thin base layer. Thethin sediment cover does not allow for the typicalhigh resolution seismic reflection profiling used todiscern landslide history on continental margins.

1.1. PreÕious slump studies on the Kilauea southflank

The south flank of Kilauea is located seaward ofits summit caldera and is bordered by the Kilauea

Ž .East Rift Zone KERZ and Kilauea Southwest RiftŽ . Ž . ŽZone KSWRZ Fig. 1 Swanson et al., 1976;

.Holcomb, 1987 . Previous studies identified the southflank of Kilauea as an active rotational landslideencompassing most of the south flank down to thepre-Hawaiian seafloor and named it the Hilina slumpŽ .Lipman et al., 1985; Moore et al., 1989 . According

Ž .to the estimates of Moore et al. 1989 , the slump isabout 5200 km2 in area and extends from the on-shore junction between the KERZ and the submarinePuna ridge to nearly the submarine continuation of

Ž .the Mauna Loa Southwest Rift Zone MLSWRZ , aregion approximately 100 km wide. In their interpre-

Ž .tation, Moore et al. 1989 show the headwall of theslump to be the KSWRZ to the northwest and possi-bly the Kaoiki fault system separating Kilauea fromMauna Loa, which is 5–10 km upslope from theHilina fault system. The KERZ continues to boundthe slump to the northeast.

Fig. 1. Shaded relief bathymetry and topography of the Hawaii island southeast flank with illumination from the SW. Box is location of Fig.3. ROV Kaiko dive 92 location illustrated by black triangle near the ridge to the SE of the base of Loihi. Dives on the Hilina slump shown

Ž .on Fig. 3a. See Fig. 3d for zone boundaries of slump. Epicenter of 1975 Kalapana earthquake 7.2 M shown by white star.

( )J.R. Smith et al.rJournal of Volcanology and Geothermal Research 94 1999 59–88 61

A more recent assessment suggests that the toe ofthe slump degrades into a debris avalanche deposit

Ž35–50 km off the present coastline Moore et al.,.1994a,b , which may originate from pre-Kilauea

landslides. The lack of detailed sampling and obser-vation has made it impossible to confirm the sourcevolcano of the debris. The recent works of Moore et

Ž .al. 1994a,b and the work of Moore and ChadwickŽ .1995 dramatically expand the lateral extent of theslump to encompass most of the southeast side ofPuna ridge. However, other recent structural studieshave limited the size of the slump to a smaller main

Žflank failure Clague and Denlinger, 1994; Denlinger.and Okubo, 1995 more in line with earlier studies

Ž .Lipman et al., 1985; Moore et al., 1989 .

1.2. Sub-aerial slump studies

The slope movement classification scheme ofŽ .Varnes 1978 , updated from his original 1958 work,

remains a classic reference on this subject. Hisscheme covers all types of movement and slopematerials. The Hilina slump best fits into the cate-

gory of rotational rock slump or translational rockblock slide; both comprise a few relatively intactunits. The mobile south flank of Kilauea also dis-plays aspects of ‘‘flows in bedrock’’ classification of

Ž .Varnes 1978 , more commonly referred to as‘‘creep’’. He notes that at the time of his writing,this type of movement in bedrock had only recentlycome under close scrutiny and was being recognizedmore frequently in high relief areas, such as moun-tain ranges in Scandinavian countries. Flows inbedrock can occur as creep, gravitational slope de-

Ž .formation, Sackung ‘‘sagging’’ , deep creep ofŽslopes, or as gravitational faulting Radbruch-Hall,

.1978 .Ž .An application of the model of Varnes 1978 to

the Kilauea south flank requires the appropriatenomenclature, shown in italics and described below,and is presented diagrammatically in Fig. 2. Themain scarp is the steep face on the undisruptedground surrounding the periphery of the slide, result-ing from the motion of the slide away from thissurface. The surface of rupture is the projection ofthis surface beneath the disturbed material. Undis-

Ž .Fig. 2. Schematic diagram of a rotational landslide with main features labeled and described in the text. After Varnes 1978 .


rupted material abutting the highest portions of themain scarp defines the crown. The head is the upperportion of the slide block along the contact betweenthe displaced material and the main scarp. The flanksare the right and left sides of the slide lookingdownslope from the crown. The main body is thedisturbed material which overlies the surface of rup-

Ž .ture where failure occurs between the main scarpŽand the toe of the surface of rupture intersection of

lower part of the surface of rupture with the original.ground surface . The toe is the perimeter of dis-

placed material farthest from the main scarp.In addition, two main zones are proposed by

Ž .Varnes 1978 . The region of depletion is the areawhere the disturbed material subsides beneath thelevel of the original ground surface. Conversely, theregion of accumulation is the area where the dis-rupted material rises above the original ground sur-face.

2. Instruments and methods

2.1. Bathymetry

Ž . ŽThe digital elevation model DEM Figs. 1 and.3a consists mainly of SEA BEAM multibeam sonar

swaths merged with nearshore hydrographic surveydata and offshore soundings from vessels of opportu-nity, along with USGS topography for Hawaii islandŽChadwick et al., 1993; Smith, 1994; Smith et al.,

.1994 . The data were processed and gridded at 0.05-Ž .min cell intervals 92 m at 198N using various

components of the freeware UNIX software pack-Žages: GMT-System Smith and Wessel, 1990; Wes-

. Žsel and Smith, 1991, 1995 , MB-System Caress and. ŽChayes, 1996 , and SBMAP Macdonald et al.,

.1992 . The entire grid was filtered using a medianboxcar filter of 0.75 km to smooth transitions be-tween different data sets and remove outliers.

2.2. Sidescan

HAWAII MR1 was used to ensonify the entiresoutheast submarine flank of Hawaii island fromnearshore to the island’s base at abyssal depths —

2 Ž . Žan area of nearly 10,000 km Fig. 3b Smith et al.,.1994 . HAWAII MR1 is a 11r12-kHz sidescan sonar

Ž .and swath bathymetry instrument Rognstad, 1992which replaced SeaMARC II. A 10-km swath widthwas used for this survey. Sidescan sonar resolution issimilar to that of SeaMARC II on the basis of aqualitative comparison of data sets, though resolutionis much better than that of the 30-km swath width

Ž .6-kHz GLORIA system Laughton, 1981 used inŽearly mapping of the Kilauea south flank Moore et

.al., 1989 . The MR1 bathymetric data were not usedin this study because complete SEA BEAM coverageof higher resolution was acquired over the same area.Lineations and slump features were determined by

Ž .overlaying sidescan and bathymetry data Fig. 3c, d .

2.3. Seismic reflection

A comprehensive single-channel seismic reflec-tion survey using a sparker source was performed

Ž . Ž .Fig. 3. Enlargement of Hilina slump area. Location in Fig. 1. a Shaded relief SEA BEAM bathymetry illumination from SW with 100 mŽ .contours. Features described in text are labeled. Numbers correspond to ROV Kaiko dives with track shown by small solid lines. bŽ .HAWAII MR1 sidescan sonar with white as high amplitude returns. Dashed bathymetric contours every y1000 m for reference. c

Ž .Structural lineation map derived primarily from sidescan and bathymetry, with verification from seismic reflection profiles Figs. 5–9 .Ž . Ž .Ticks on face where highest acoustic reflectivity could be determined, indicating down-dropped side. Interpretations in c and d are

Ž . Ž . Ž .overlayed on shaded bathymetry. Epicenter of 1975 Kalapana earthquake shown by white star in c and d . d Geologic interpretation mapŽ .with Zones 1–4 as labeled. Same data as c . Dotted pattern nearshore is hyaloclastite zone. Undulating outlines with shaded patterns at baseŽ .of Kilauea and Loihi are basal lava flows Blacksgreater reflectivity; Grays lesser reflectivity . Thick solid lines separate major geologic

features. Thin solid lines denote major Hilina slump zones offshore and basal blocks. Thick dashed lines show less resolved or extrapolatedportions of slide boundaries on outer scarp and continue downslope, encapsulating debris fieldrtoe area. Thin dashed lines are contacts

Žwhere resolution is poor. Dash–dot lines outline bathymetric benches. Cross-hatched lines are Loihi rift zones with dashed ones as.proposed extensions .



Ž .Fig. 3 continued .


aboard RrV Morskoy Geofizik. More than 3600 kmŽ . 2of profiles Fig. 4 over nearly a 20,000-km area

Ž .were collected Smith, 1996 . A ;100-m seismichydrophone streamer was towed ;200 m behindthe vessel and a ;5-kJ sparker source was triggeredabout every 15 s. Lines were spaced ;3.7 km apartover the main flank to the base of the island, then

Ž;7.4 km thereafter out to ;130 km offshore Fig..4 . The profile data are presented in Figs. 5a, 6a, 7a,

8a, 9a and line drawing interpretations are shown inFigs. 5b, 6b, 7b, 8b and 9b.

2.4. ObserÕations and sampling

Detailed sampling and observations of the sus-pected lateral boundaries, outer scarp, and downs-lope debris were carried out for the first time inSeptember 1998 using the ROV Kaiko. This allowed

Ž .us JRS on scientific team to determine the compo-sition of the lower slope material. Four dives were

Ž .made on the Hilina slump Fig. 3a and one near theŽ .basal ridgerblock to the southeast of Loihi Fig. 1 .

From a tether to its mothership, RrV Kairei, the

Ž . Ž . Ž .Fig. 4. Numbers and thick lines show location of seismic reflection profiles in Figs. 5–9. A Tracks for Figs. 5 and 7 labeled . B TracksŽ .for Fig. 6. C Tracks for Fig. 8. Vectors and error ellipses show average horizontal velocities determined from GPS surveys between 1990Ž .and 1993 after Owen et al., 1995 . Length of longest vector represents ;10 cmryear and ellipses depict 95% confidence intervals.

Ž .Epicenter of 1975 Kalapana earthquake shown by white star. D Tracks for Fig. 9.


Fig. 5. Marine single-channel seismic reflection profiles as located in Fig. 4A. Zone 1, the upper slope of the upper main body of the HilinaŽ .slump. a Data profiles. Time in seconds is two-way travel time. Horizontal gradationss1r2 s two-way travel time. Vertical

Ž .gradationss30 min ship travel time. Vertical exaggeration ;14= . b Line drawing interpretation.


Ž .Fig. 5 continued .

robot submersible can descend to a maximum depthof 11,000 m. Kaiko has video and still cameras for

observation, and two manipulator arms and a samplebasket for collecting rocks and sediment push cores.Both vessels are owned and operated by the Japan

Ž .Marine Science and Technology Center JAMSTEC ,and this science program was a collaboration be-tween Japan, University of Hawaii, and USGS scien-tists.

3. Hilina slump geomorphology

We have divided the offshore portions of theHilina slump described in this paper into subsections

Ž .corresponding to the nomenclature of Varnes 1978for a rock slump or rock block slide, simply referredto here as a rotational landslide. It is important tonote the designation of the lateral boundaries, herereferred to as the right and left sides rather thanflanks to avoid confusion with reference to the wholevolcanic flank. The regions of depletion and accu-

Ž . Ž .mulation Varnes calls them zones of Varnes 1978are further subdivided, for the Hilina Slump, intofour zones, two within each region. Zones are re-ferred to in Fig. 3d and Figs. 5–10 and datarinter-pretations are presented in Figs. 1, 3 and 5–10. Theslump zones should not be confused with establishednames for the several sub-aerial fault zones in thesefigures.

3.1. Region of depletion

3.1.1. Zone 1: Upper main body — onshore to thenearshore slope

The crown is a sub-aerial feature corresponding tothe Kilauea rift zones and Kaoiki fault system, whilethe head, main and minor scarps form the Koae andHilina fault systems. These features are thoroughly

Ždescribed in other works e.g., Moore and Krivoy,.1964; Holcomb, 1987 and summarized here within

the context of a slump.

3.1.1.1. Crown. The KERZ is 4–6 km wide andextends 60 km from the caldera to the east cape of

ŽHawaii island into the submarine environment Fig..1; see below . The rift zone is divided into an upper

segment with abundant pit craters, and middle andlower segments showing closely spaced fissures,faults, and grabens with ‘‘net seaward subsidence’’


Ž .Holcomb, 1987 . The axis of the KERZ seismicityband is displaced approximately 0.5 km south of thesurface expression of the rift as defined by the

Žvolcanic features mentioned above Klein et al.,.1987 .

In comparison to the KERZ, the KSWRZ doesnot form an obvious ridge and its seaward subsi-

Ž .dence is more pronounced Holcomb, 1987 . TheKSWRZ includes an inactive zone inland and anactive zone seaward dividing the rift zone into a

Ž .lower and an upper segment Holcomb, 1987 , nearlyopposite to the activity pattern of the KERZ. The

Ž .active seaward lower segment is characterized by‘‘strands of closely spaced fractures separated by

Žunbroken blocks’’ and is about 1 km wide Holcomb,.1987 . The majority of faults are down-thrown to the

southeast with as much as 20 m of displacement.Ž .Holcomb 1987 notes that a seaward migration of

volcanic activity may have taken place since alleruptive vents seaward of the faults appear youngerthan the inland vents which are greater than 1000-years-old.

Like the KERZ seismicity, the KSWRZ earth-quake pattern shows a narrow and almost continuousband of activity with the primary magma conduit

Ž .located at about 3 km depth Klein et al., 1987 .Similar to the KERZ, the seismicity band divergesfrom the surficial expression of the rift up to 2 kmsouth. As with the Kilauea caldera offset, the dis-placement of seismicity from surface geology could

Žresult from southward migration of the rifts Swan-.son et al., 1976 .

The northeast-trending Kaoiki–Waiohinu faultŽ .system of Mauna Loa Fig. 1 is thought to be a

relatively inactive zone of slump block structures,similar to the more active Hilina fault system, formedby seaward-directed gravitational slumping of the

Ž .Mauna Loa southeast flank Lipman et al., 1990 .However, the epicenter of the 1982 Hilea earthquakeŽ .Ms5.4 was localized at the Kaoiki–Waiohinu

Ž .fault system Wyss and Zhengxiang, 1989 . Earth-quakes along the Kaoiki–Waiohinu system supportthe gravitational slump model proposed for the

Žsoutheast slopes of Mauna Loa Endo, 1985; Wyss,.1988 .

3.1.1.2. Head, main and minor scarps. The HilinaPali faults, a prominent system of normal faults on

the south flank, are down-thrown to the south withindividual scarps up to 500 m high. The net throw onthe entire Hilina fault system is approximately 600m. They are mantled in many places by seaward-flowing lavas from recent and current volcanic erup-tions of Kilauea and the KERZ. The most prominent

Žfaults are located in the seaward zone Swanson et.al., 1976; Holcomb, 1987 . No lavas are known to

have erupted along the Hilina fault system, and onlya few dikes have been mapped on the southeast flankŽ .Walker, 1969; Easton and Lockwood, 1983 . How-ever, displacement along the Hilina fault system isbelieved to dilate the rift zones and disturb the

Žmagmatic plumbing of Kilauea Swanson et al., 1976;.Holcomb, 1987; Klein et al., 1987 .

From the presence of anomalously steep slopes,Ž .Holcomb 1980 suggests that buried Hilina faults

exist inland near the Koae fault system and arepossibly found in the KERZ and KSWRZ. Recentmovements along the Hilina Pali indicate the faultsystem is isolated from other Kilauea structures, but

Ž .Holcomb 1987 proposes that the Hilina fault sys-tem ‘‘grades into’’ other features. He concludes thatthe age of the Hilina fault system might pre-dateKilauea, developing from seaward migration of the

Ž .Kaoki fault system on Mauna Loa as older inlandfaults were stabilized by the growing Kilauea edifice.

Maximum vertical movement of 3.5 m with 8 mof horizontal tensional displacement across the Hilinafault system were reported in 1975, coincident withthe magnitude 7.2 earthquake located below Kala-

Žpana on the southeast coast of Kilauea Lipman et.al., 1985 . They showed that the boundary between

domains of extensional and compressional strain isclearly submarine and may be located relatively lowon the south flank. Both the vertical and horizontalsub-aerial displacements increase from the summittowards the coastline, but no data were availablebelow sea level to define the boundary more pre-

Ž .cisely. The study by Lipman et al. 1985 , along withŽ .that of Swanson et al. 1976 , suggested that a

recurrent pattern of dike injection followed bydownslope, seaward failure on the Hilina Pali is partof the normal volcanic growth process of Kilauea.

Sonar images show a major portion of the uppersubmarine slope of the south flank is dominated bydownslope-oriented streaks of relatively high reflec-tivity, which are interpreted as surface mantling of


the slopes by downslope flows of hyaloclastite andŽ .lava Fig. 3b . Seismic profiles also indicate hyalo-

Ž .clastite layering Fig. 5 . These materials result fromlava delta collapse with sub-aerially erupting lavafragmenting when it reaches the shoreline. The de-posits blanket most of the geologic features down to;2500 m water depth with a relatively featurelessmask of coarse sediment.

The onshore Pali Uli, intercepting the coastline atŽ .y1558 near Kapaahu Fig. 1 , lies along the same

east–west-striking trend as a pair of 5-km-long lin-eations, possibly longitudinal faults, on the lowerreaches of the upper slope centered at 19818X,

X Ž .y154853 , 2500 m water depth Fig. 3b,c . Theselineations intersect the normal offset faults evident

Ž .on the northeast edge of the outer scarp Fig. 3d .Though much of upslope evidence is masked by thenearshore lava flow debris, these lineations may beevolving fault systems, perhaps connected at depthand partially defining the northeast lateral boundaryof the slump. Other workers have placed the north-

Žeast slump boundary near this location Lipman etal., 1985; Clague and Denlinger, 1994; Denlinger

.and Okubo, 1995 .

3.1.2. Zone 2: Lower main body — midslope benchand catchment basin

A catchment basin lies on the northeast portion ofŽ w xthe midslope bench Figs. 3a,d, 6 profiles 2, 3 and

.Fig. 7 and is most likely bounded by active faults,owing to the reorientation of the sediment layers

Ž .through tilting or down-bowing Fig. 6 . In addition,the extensional nature of the upslope region is well-

Ždocumented Lipman et al., 1985; Delaney et al.,.1990; Owen et al., 1995 . The basin extends 15 km

parallel to slope, and ranges from 2 to 10 km wideand 2600–2900 m water depth. It covers an area ofapproximately 60 km2, with a sediment thickness upto 0.25 s two-way travel time from seismic reflection

Ž .records Fig. 6 . A conservative sound velocity of1800 mrs for the material provides a 180–240 msediment thickness, which computes to an existingsediment volume of 11–14.5 km3. Stratified layers

Ž .are generally oriented horizontally Fig. 6; profile 2Ž .to sub-horizontally Fig. 6; profiles 1, 3, 4, 6 in the

basin. With the estimate of the volume of materialsupplied to the shoreline averaging 500,000 m3rday

Ž .HVO Staff, 1994 , it would take 120–170 years ofcontinuous flow to accumulate this much material inthe basin at a rate of ;0.1 km3ryear. This rate ishalf that of the lava reaching the coastline andflowing into the sea because we know that the upper

Žslope is mantled by volcaniclastic sand Lee et al.,.1994 . Such accumulation would take much longer

because, historically, lava has not entered the sea forsuch continuous time periods.

A ridge with 200–400 m relief, and a widthranging from 0.5 to 7 km, is located on the bench

Ž waxis, south of the catchment basin Figs. 3a, 6 pro-x .files 4, 5 and Fig. 7 . Two enclosed circular depres-

sions, 400 and 200 m deep and 5.5 and )2 kmwide, respectively, are located along the prevailing

Ž w x .508 strike azimuth Figs. 3a, 6 profile 6 and Fig. 7 .They alternate in position with portions of the ridge.Sediments on the 400-m-high part of the ridge,somewhat dome-like in shape, are approximately thesame thickness as those in the catchment basin. This

Žridge, though backing the transition zone as Zone 3,.described in Section 3.2.1 , appears distinct from

Žthat structure and the midslope basin Fig. 6; profiles.4, 5 . While the ridge is generally heavily sedi-

mented, the seaward side is highly reflective in theŽHAWAII MR1 sidescan data Fig. 3b, see Fig. 3a for

.reference , similar to a lava flow, volcanic mound, ora scarp face. Floors of the depressions are sedimentcovered, but their inner walls are highly reflective,indicating steep slopes and possibly active collapse.The fault-bounded spire on the seaward edge of the

Ž .midslope bench Fig. 6; profile 5 is a curious struc-ture and may imply more than just uplift of sedi-ments due to compression of the lower flank. Itcould be an outcrop of upthrust indurated sediment

Žfrom below Zones 3 and 4, described in Sections.3.2.1 and 3.2.2 .

3.2. Region of accumulation

3.2.1. Zone 3: Transition — transÕerse cracks andtoe of surface rupture

A faulted andror fissured zone 18-km long and3–7-km wide, trending 508, cuts across the seawardedge of the midslope bench subparallel to the pre-

Ž .vailing slope Figs. 1, 3c and 6 . It forms a lip with100–200 m relief near the seaward edge of Zone 2 at



Fig. 6. Marine single-channel seismic reflection profiles as located in Fig. 4B. Zones 2 and 3, the midslope bench and transition zone,Ž .respectively, of the lower main body of the Hilina slump. a Data profiles. Time in seconds is two-way travel time. Horizontal

Ž .gradationss1r2 s two-way travel time. Vertical gradationss30 min ship travel time. Vertical exaggeration ;14= . b Line drawinginterpretation.

2500 m, which traps the upper slope volcaniclasticmaterial, preventing its movement farther downs-

lope. The lower boundary of this disruptive band oflineations marks the slope break to the outer scarp at


Ž .the 2800 m isobath Fig. 3a–c . The fractured zoneis composed of lineations 1–3-km long with 10–40

Ž .m relief some up to 150 m relief following theŽ .same orientation as the major trend Fig. 3c . This

band of lineations resembles the transverse crackzone characteristic above the toe of rupture surface

Ž .of the Varnes rotational landslide model Fig. 2 . Itcould also represent an imbricate fault or nappe

Žstructure characteristic of a thrusting regime De-.nlinger and Okubo, 1995 . Heavy sedimentation is

evident shoreward of the upper lip of this transitionzone. Contrarily, the seaward face appears highly

Žreflective and disrupted Fig. 3b, below the y3000w x.m contour see Fig. 3a for reference , suggesting a

sheared off failure surface owing to the numerousŽ .lineations visible in the sonar imagery Fig. 3b,c

rather than simply being sediment-starved.

3.2.2. Zone 4: Foot

3.2.2.1. Upper foot — outer scarpBelow the fractured transition zone, the outer

scarp begins at a water depth of 2800 m and de-scends to approximately 4700 m with an average

Ž .slope of 168 Figs. 1 and 3d . Four or five smalllandslides with concave headwalls and downslope

Ž .debris fields incise this part of the flank Fig. 3a–d .We believe that these small erosive events weretriggered by the deformation of the seaward face ofthe transition zone caused by the impinging Kilaueavolcano flank. The sidescan imagery shows chaotic

Ž .patterns of high and low reflectivity Fig. 3b and theseismic reflection profiles show almost no surface

Ž .return andror side reflections hyperbolae as a re-sult of the rapid drop-off and abundant hummocky

Ž .debris Fig. 8 . These observations are consistentwith the region having undergone one or more

Žepisodes of failure Lipman et al., 1988; Moore et.al., 1989 .

From observations and samples taken during theROV Kaiko dives, it was determined that above;3000 m the lower portion of the upper slope iscomposed of pillow breccia from Kilauea as ex-

Ž .pected Fig. 3a; dive 95 . However, below 3000 mŽ .on the outer scarp Fig. 3a; dives 91 and 98 , pillows

and pillow-breccia were absent, replaced instead bymassive outcrops of sandstone–mudstone composed

Žof indurated volcanic sediments Naka et al., 1998;

.Shipboard Scientific Party, 1998 . Preliminary analy-ses of the glass chemistry suggest there are noKilauea components or Mauna Loa tholeiitic sourcesrepresented in this 2-km-thick section. Instead, itappears that Mauna Kea andror Kohala are the

Ž .source volcano s for the material contained in theseŽ .sandstones Lipman et al., submitted . Whatever the

composition, the outer scarp is a pre-Kilauea featureand represents an unstable zone undergoing possiblyrecurrent mass wasting. The sediments have likelybeen scraped off the pre-existing slope by the ad-vancing Kilauea flank, indurated by compression,and piled into a thickened wedge subject to gravita-tional failure from oversteepening and seismic trig-gering.

3.2.2.2. Lower foot — blocks and toesDetached blocks. A debris field with an area of

about 600 km2 extends seaward from the base of theŽ .outer scarp Figs. 1 and 3a,d . Some 25–30 detached

blocks, rounded or elongate in shape, are presentacross this 30–35-km-long stretch of the lower slopevarying between 8 and 20-km wide below the outerscarp at 4700–5200 m water depth and 35–50 km

Ž .offshore Figs. 3a–d and 9 . The larger elongateblocks are up to hundreds of meters in relief, kilome-ters wide, and up to 10 km in length. The largest ofthese blocks was studied during the ROV Kaiko

Ž .cruise Fig. 3a; dive 93 : this block, and likely all theothers, is composed of the massive sandstones de-

Žscribed in Section 3.2.2.1 Naka et al., 1998; Ship-board Scientific Party, 1998; Lipman et al., submit-

.ted . A loosely consolidated mudstone was recoveredfrom the basal ridgerblock to the southeast of Loihi

Ž .at ;4750 m water depth Fig. 1; dive 92 .Sediments folded into an anticline between the

outer scarp and the large ridge-like block at 198N,X Ž w x.154850 W Figs. 3a, 8 profile 2 suggest compres-

sion, perhaps from the creeping upper foot of theslump. A bulge in the bathymetry at 19802X, y154851X

Ž .Fig. 3a suggests that this accumulating mound ofsediment is being rerouted and flows to the northeastas a result of the damming effects of the large block.The seafloor on the seaward side of this large block

Ž .is 300 m lower than the landward side Fig. 3a .Several adjacent blocks near the western edge of theridge-like block appear as lineations in the HAWAII

Ž .MR1 sidescan data Fig. 3b, c , forming a 2-km-wide

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Ž .Fig. 7. Marine single-channel seismic reflection profiles as located in Fig. 4A–C. Along strike and across Zone 2, the midslope bench and catchment basin Fig. 6 of the HilinaŽ .slump. NE and SW lateral boundaries of the slump are shown. Interpolated dashed line where weak returns caused data record to fade. A Data profiles. Time in seconds is

Ž .two-way travel time. Horizontal gradationss1r2 s two-way travel time. Vertical gradationss30 min ship travel time. Vertical exaggeration ;14=. B Line drawinginterpretation.



Fig. 8. Marine single-channel seismic reflection profiles as locatedin Fig. 4C. Zone 4, the outer scarp of the upper foot of the Hilinaslump. Two profiles over the remnant hilina slump block are

Ž .included see Fig. 3d, 10D . Interpolated dashed line where weakŽ .returns caused data record to fade. a Data profiles. Time in

seconds is two-way travel time. Horizontal gradationss1r2 stwo-way travel time. Vertical gradationss30 min ship travel

Ž .time. Vertical exaggeration ;14=. b Line drawing interpreta-tion.

and 5-km-long chute that may provide a major path-way for terrestrial sediment to reach the HawaiianDeep. The Hawaiian Deep is the most southern anddeepest extent of the Hawaiian moat, a flexuraltrough surrounding the southeastern Hawaiian Ridge

Žwith a radius of approximately 140 km Hamilton,.1957 .

Thrust toe and transÕerse ridges. Southeastwardof the suspected lava flows and the detached blockzone at 33–53 km from shore, there occur offsets inthe sediments with relief up to 75 m. These resemble

Ž .the transverse ridges in the model of Varnes 1978Ž .Fig. 2 and suggest that the toe of the slump ismanifested as a series of compressional ridges. Theridges are discernible on adjacent seismic reflection

Ž .profiles Fig. 9 as possibly being uplifted above theŽsurrounding seafloor. The first set of ridges closest

.to shore are more pronounced and contiguous be-Ž .tween seismic reflection profiles Fig. 9 . These data

are used as the basis for the dashed line in Fig. 3dŽ .Toe . They are located just seaward of the debrisfield that includes the large elongate blocks previ-ously defined as the toe of the slump based upon

Ž . Žmodels of 1 thrust faulting Borgia et al., 1990;. Ž . ŽBorgia and Treves, 1992 , 2 delta formation De-. Ž .nlinger and Okubo, 1995 , and 3 low-resolution

Ž .bathymetry Lipman et al., 1985 .A second series of disruptions in the sediment,

though far less prominent, emerge 10–20 km fartherŽ .offshore Fig. 9 from the first set. Though this

second series of offsets corresponds nearly to theŽ .Moore et al. 1989 slump toe, it is not prominent

enough to reliably define it as the toe and distinguishit from structures possibly related to the HawaiianDeep. In addition, the same secondary features arenot seen on every profile. On the profiles in whichthey are evident, identical uplifts also occur nearly130 km offshore — near the end of our seismicreflection tracklines. The uplift structures at the sec-ondary toe resemble those proposed on the basis of

Ž .seismic reflection data by Normark and Shor 1968Ž .their fig. 7 as resulting from tensional fracturing ofthe Hawaiian Arch. In summary, the toe, as defined

Ž .in this paper Fig. 3d , is delineated by the first set ofuplifts and forms an arc 35–40 km wide. Data on thenortheast and southwest edges of the feature are lessreliable. A summary of the data in Figs. 3 and 5 ispresented in Fig. 10 and is discussed in Section 4.


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Fig. 9. Marine single-channel seismic reflection profiles as located in Fig. 4D. Zone 4, the toe area of the lower foot of the Hilina slump. Primary toe is distal end of slump.Ž .Interpolated dashed line where weak returns caused data record to fade. a Data profiles. Time in seconds is two-way travel time. Horizontal gradationss1r2 s two-way travel

Ž .time. Vertical gradationss30 min ship travel time. Vertical exaggeration ;14=. b Line drawing interpretation.



4. Discussion

It is known from previous studies that abruptmovement of the south flank occurs coseismically

Ž .with major earthquakes e.g., 1868, 1975 releasingdecades of stress build-up from dike intrusion intothe rift zones. The stress was created by a lateralcompressive stress along the buttressed north side

Žand a less-supported south flank Swanson et al.,.1976; Lipman et al., 1985 . Thus, ‘‘magma-jacking’’

is at least partially responsible for the gradual devel-opment of the headwall and lateral boundaries byincremental and repeated pressure over time causingepisodic seaward displacement of the rift zone and

Ž .flank by the creep process. Duffield et al. 1982 , intheir comparison of Piton de la Fournaise and Ki-lauea, suggest that periods when magma pressure iselevated in the reservoir system with concurrent

Ž .injection of magma into the rift zone s would pro-vide prime conditions for ‘‘detachment and slumpingof a rift zone- and hyaloclastite-bounded block.’’

Since Kilauea is built on the formerly unbut-tressed south flank of Mauna Loa and is proposed to

Žfail down to the Cretaceous seafloor Lipman et al.,.1985; Moore et al., 1989 , it must have incorporated

part of the Mauna Loa slope into its sliding slumpmass. This may be one of the determining factors inthe development of a slump vs. a debris avalanche:growth upon the flank of an adjacent older volcanicedifice. This is evident when one considers that otherknown slumps on the southern Hawaiian Ridge are,in most cases, at least the second volcano to develop

against an existing edifice, as seen with the Hana,ŽNorth and South Kona, and the Hilina slumps Moore

.et al., 1989 . Though the available examples are notgreat in number, only the Waianae slump, from the

Žolder Waianae volcano i.e., older than Koolau vol-.cano on Oahu, varies from this trend. The primary

slope may provide the unstable slip surface for thesecondarily developing volcano, especially when oneconsiders the recent discovery of a thickened andhardened sedimentary wedge beneath the south flank

Žof Kilauea Naka et al., 1998; Shipboard Scientific.Party, 1998 .

As we will discuss in the following text, it isincorrect to refer to the mobile south flank of Ki-lauea as ‘‘unbuttressed.’’ Though it is in motionseaward and not stabilized by a large edifice abovesea level, the three seamounts at the base of, or onthe southeast flank itself including Apuupuuseamount to the southwest, appear to be greatlymodifying downslope flank deformation. Withouttheir presence, the flank might move more rapidly.Referring to the south flank as the ‘‘less-buttressed’’flank or simply the ‘‘mobile’’ south flank is moreappropriate.

To arrive at the current stage of development ofŽ .the submarine south flank Fig. 3d , a morpho-struct-

ural evolutionary model for Kilauea volcano and theŽ .Hilina slump is proposed Fig. 10 . This model does

not attempt to determine the location of the paleo-shoreline for each successive step, which would beaffected by eustatic sea level and local subsidence.The present day coastline is illustrated in Fig. 10A–C

Fig. 10. Evolutionary model for Kilauea volcano and the south flank highlighting development of the Hilina slump. This model does notattempt to account for position of paleo-shoreline, instead, it considers development of the entire edifice without regard to sea level. Arrows

Ž . Ž .in this figure only indicate direction of motion and are not vectors. Details described in the text. Modified after Smith 1996 . A At ;90Žka. Heavy dash–dot lines are rift Kilauea rift zones. Thin solid lines are developing fault systems, fractures, and fissures except for confetti

.pattern on white background . Light dotted line is present shoreline for reference. Thrusting indicated by triangles on feature boundary.Medium-sized black arrows indicate direction of flank growth and motion. MLsMauna Loa, KsKilauea, LsLoihi submarine volcano,

Ž .AsApuupuu seamount, HsHohonu seamount. B At ;60 ka. Large black arrows indicate proposed directions of motion of centralsouth flank and developing Hilina slump. Small black arrow shows direction of motion of developing SW boundary adjacent to Papau ridgeŽ . Ž .P , depicted as curved, overlapping en echelon structural features same on NE boundary . Smallest arrows indicate direction of opening of

Ž .transition zone and of avalanching debris from the outer scarp see D . Ticks on normal faults indicate down-dropped side. Open trianglesŽ .indicate thrusts, some with dashed boundary, where faulting and boundaries are suspect. C At ;30 ka. Large white arrows in place of

Ž . Ž .black ones from B , for reference, show maturation to a counterclockwise rotational fall-away downslope motion heavy dashed arrow , asŽ Ž .. Ž .depicted in the inset. Additional small arrows on SW slump boundary and NW Loihi illustrate fault motion also in D . D Now. Fully

developed rotational motion illustrated by heavy solid arrow. Principal flank and slump features labeled. Transition zone is lineated areaŽ .between basin on bench and outer scarp see Figs. 1 and 3d . Present coastline shown in heavier dotted line than A–C. Proposed slump toes

in light dashed and dotted lines. Kilauea–Mauna Loa sub-aerial boundary traced by heavy solid line continuing offshore.


as a dotted line only for reference. In a paper on thegrowth and evolution of Hawaii island, Moore and

Ž .Clague 1992 depict a young Kilauea volcano at100 ka with no flank failures. In keeping with theirmodel, the evolution proposed below begins nomi-nally at 90 ka.

In the model, the young Kilauea volcano southŽ .flank, with a typical convex insular slope Fig. 10A ,

extends from the KSWRZ to the KERZ near the easttip of the island and eventually beyond as the Punaridge develops seaward. Modern day fault systemsŽ .Kaoiki, Koae, Hilina begin to develop early on andevolve gradually over time in response to the grow-ing mass of the Kilauea edifice. Loihi seamount

Žeither does not yet exist or is still relatively small as.shown and does not yet affect the moving flank. An

upper delta of terrestrially derived and coastal vol-canogenic material begins accumulating downslopeof Kilauea caldera. The Kilauea south flank spreadsseaward, overthrusting the previously deposited

Žglass-rich sands Naka et al., 1998; Shipboard Scien-.tific Party, 1998 which act as a decollement. The´

sediments are thickened into a wedge by compres-Ž .sion Lipman et al., submitted as the Kilauea south

flank plows its way downslope. The compressiveforces cause the sediments to become indurated and

Ž .continue to build a higher outer frontal scarp. Suchfeatures have been described resulting from large

Ž .sub-aerial landslides Varnes, 1958 and volcanoesŽ .van Bemmelen, 1949; Borgia et al., 1990 . Mean-while, the thrusting motion and associated earth-quakes, combined with a steepened outer scarp, leadsto massive blocks breaking off and forming thedebris apron below.

Loihi continues to grow and begins buttressingthe impinging flank and halts its movement southeastuntil a new echelon fault system begins to developŽ . Ž .Fig. 10B . Fleming and Johnson 1989 report thatlarge landslides are laterally bounded by structuresanalogous to strike-slip faults. The scarp in the mid-dle of the southeast flank associated with Papauridge may indeed represent such a fault structure andserve as the present right-lateral boundary of theHilina slump, opposite Loihi seamount. Day et al.Ž .1997 report a similar lateral slump boundary on theCanary Island of El Hierro. The lineations inter-preted to be a series of 2–5-km-long transverse

Ž .fractures in the re-entrant east of Loihi Fig. 3c

suggest that a type of strike-slip or tearing displace-ment is located there. This leaves the southwestportion of the slump separated from the main drivingforce, which is seaward of the main caldera andKERZ plumbing systems. It continues to grow anddownstep to the southeast until the KSWRZ becomesrelatively inactive in a later stage, possibly related tothe readjustment of the local stress regime andror aredirected magma plumbing system following theinitiation of flank displacement.

At this stage, Papau ridge begins to develop,possibly as a fold in response to compressional

Ž .forces Morgan et al., 1998 combined with shearingat the new block boundary. This development resultsfrom the summation of forces that eventually pro-

Ž .duce rotational motion Fig. 10B . The sidescansonar data presently show lineations along the topand southwest side of Papau ridge oblique to the

Ž .direction of motion Fig. 3b, c . These structuresmay be interpreted as caused by tectonic shearing.Another hypothesis is that Papau is a structuralcompression ridge accommodating convergence be-tween the Hilina slump and the stable ground surfaceŽ .Moore and Chadwick, 1995 . At this stage, thedisplacement mode of the south flank may take on

Žaspects of volcano spreading Borgia and Treves,.1992 , defined as extensional faulting at the volcano

summit coincident with compression structures at theŽ .base Borgia, 1994 . Conditions required for vol-

canic spreading are the existence of a weak basallayer, an adequately large volcanic and intrusivecomplex, and an ample magma input rate to drive theprocess. Five nominal stages define this process:building of the initial volcanic pedestal, compressioncaused by loading, thrusting along basal decolle-´ments, intrusion of volcanic complexes, and spread-

Ž .ing of the edifice and complexes Borgia, 1994 .Almost the entire south flank may still be sliding

Ž .seaward Fig. 10B , extending to a pivot point fur-ther to the northeast along Puna ridge. A rotationalpole located on Puna ridge satisfies some modelsproposed in the literature where the Hilina slumpincludes nearly the entire submarine KERZ to

X Ž154820 W e.g., Clague et al., 1994; Moore et al.,.1994a,b; Moore and Chadwick, 1995 . Clague et al.

Ž .1994 note the existence of a series of broad basalbenches and linear ridges or terraces along the south-east side of the Puna ridge that they interpret as a


Ž .series of rotational slump blocks Fig. 1 . Theyconclude that the structures of the mobile south flankand rift zone axis are linked. Moore and ChadwickŽ .1995 suggest that these benches and blocks formthe eastern slump boundary. Volcanic constructioncontinues.

Ž .In the next stage of development Fig. 10C , moreof the present day slump features start to form andthe mobile portion of the south flank may becomemore wedge-shaped in plan view. Loihi grows largeras the southwest slump boundary matures from itsechelon beginnings. The transition zone continues todevelop, marking the advent of sufficient compres-sive forces, perhaps from the buttressing seamounts,to counteract some of the extension of the flankseaward. The transition zone, corresponding to the

Ž .toe of surface rupture in Varnes’ model Fig. 2 , iswhere one expects extensive buckling and fracturingto occur because of the interception of the lowerportion of the surface of rupture with the originalground surface. This zone marks the change fromupslope material depletion to downslope accumula-tion over a basal high point.

Additionally, the outer scarp thickens and be-comes more pronounced due to the advancing Ki-lauea edifice, with additional debris spalling or calv-ing off, enlarging the existing concave landslidescars. The material is deposited in the basal debrisfield at the toe of the slump. Much larger breakawayslide blocks from the Nuuanu and Wailau debris

Žavalanches Moore, 1964; Moore et al., 1989; Nor-.mark et al., 1993; Smith et al., 1998 and the South

Ž .Kona slump Moore et al., 1995 are proposed tohave been emplaced even farther from the sub-aerialflanks of their respective source volcanoes.

Seaward motion now slows as a result of thenortheast side of the slump encountering the buttressof Hohonu seamount, a feature thought to be of

Ž .Cretaceous-age Moore and Fiske, 1969 . This sideof the slump is then obstructed, allowing the south-west side to pivot farther seaward, until the new faultsystem can mature. This causes the Hilina slump totake on a counterclockwise rotational ‘‘fall-away’’motion downslope as depicted in the inset of Fig.10c. Paleomagnetic studies of outcropping lava flows

Žin scarps of the Hilina fault system Riley et al.,.1999-this volume indicate that the more seaward

Ž .fault block Puu Kapukapu has rotated 14.88"8.58

counterclockwise with respect to the block fartherŽ .inland Keana Bihopa . A remnant Hilina slump

block to the east of the newly developing northeastboundary is stalled from further downslope motion

Žand left to remain relatively stationary Figs. 3d, 7.and 10D , being fully buttressed by Hohonu

seamount. It maintains its outer scarp stepped topog-Ž .raphy Fig. 8 similar to the KSWRZ submarine

Ž . Žextension Fig. 1 Smith, 1996; Smith et al., in.preparation , described by some workers as a sepa-

Ž .rate slide unit Lipman et al., 1990 . The series ofbasal ridges at the foot of the southeast side of Puna

Ž . X Xridge Fig. 1 from longitude y154845 to y154825Ž .Clague et al., 1994 appear to bow convexly aroundHohonu, especially its northeast side, suggesting but-tressing of the southwest Puna ridge flank by Ho-honu seamount.

Further expansion of the key parts of the activeŽHilina slump continues into the next stage Fig.

.10D . At this step, the transition zone grows widerand more small landslides occur on the outer scarp,depositing additional blocks at the base. The foot andtoe areas more fully develop, with seaward sediment

Ž .disruptions evident beyond the debris field Fig. 9 .The northeast fault boundary continues to maturewhile still pivoting around Hohonu. The pivot nowshifts from farther northeast along Puna ridge to justlandward of Hohonu on the upper slope and outerscarp.

Hohonu seamount may serve both as a pivot pinand as an abutment. A saw-toothed conjugate faultpattern consisting of several step faults begins atapproximately 3000 m water depth directly landwardof Hohonu and steps down in an offset to the south-southwest where the faults intersect the base of the

Žouter scarp at 4600 m water depth Figs. 3c and.10D . As a terrestrial analogue, similar serrated pat-

terns occur near slump toes in the sub-aerial Alani–Paty landslide complex in Oahu’s sediment-filled

Ž .Manoa Valley Baum and Fleming, 1991 . FlemingŽ . Ž .and Johnson 1989 p. 61 state that echelon tension

cracks, or tear faults on a larger scale, can beassociated and interspersed with small thrusts onlandslide boundaries and may represent a strike-sliptype fault at depth which has not yet propagated tothe surface. They add that one side may be undergo-

Žing fully developed mode II deformation fracturesurfaces slip at right angles to fracture front; Lawn


.and Wilshaw, 1975 while the opposite boundary isŽexperiencing mode III deformation fracture surfaces

.slip parallel to fracture front and is surficially mani-fested as echelon faults and thrusts. These observa-tions lend support to the structural interpretation anddeformation model proposed in this study based onthe highly resolved geomorphology of the submarinesoutheast flank.

Further evidence for Hohonu acting as a pivot pinis found by extrapolating the sub-aerial motion vec-

Žtors acquired from leveling and GPS studies Lipman.et al., 1985; Delaney et al., 1993; Owen et al., 1995 .

These data display up to four times more seawarddisplacement after the 1975 earthquake and up totwo times more for a sustained eruptive period thanmeasurements in areas farther to the northeast near

Ž .Kalapana Fig. 4C . This differential of displacementmust be accommodated farther downslope. The ac-commodation zone is proposed to be the large wedgeof material forming the midslope bench and outer

Žscarp. Some researchers Moore et al., 1989; Moore.and Clague, 1992 estimate the slump is moving

seaward at a rate of approximately 10 cmryear,based on extrapolation from the motions associatedwith the 1975 event. This range of movement iscomparable to the sub-aerial creep rates from geode-

Ž . Žtic work Fig. 4C Delaney et al., 1993; Owen et al.,.1995 .

While a distinct scarp has formed on the south-west margin, perhaps because the area directlydownslope from the caldera and primary plumbingzone is prone to most displacement, the northeastlateral boundary has not fully developed and is notconnected to the basal transverse ridges. The bound-aries based on the offshore geomorphology matchwell with structural and earthquake modeling by

Ž .Denlinger and Okubo 1995 . Magma-jacking maycause the moveable portion of the flank to slowlyand quasi-continuously advance along the developed

Ž .right southwest lateral bound, all the while leadingto stress in the undeveloped left lateral boundaryŽ .northeast . This action results in a counterclockwiserotation of the slump block through the failure win-

Ždow between Loihi and Hohonu seamounts Fig..10D . Once developed, there are at least two possi-

Ž .bilities: 1 little additional stress will build up in themain Hilina slump block and thus creep will occursemi-continuously along both matured boundaries in

Ž .lieu of inducing large earthquakes; andror 2 thelarge basal earthquakes will continue to be triggered,possibly more frequently, because of the more easilyinduced slip due to insufficient friction along thenow fully evolved left lateral boundary.

Intraplate volcanic islands present a unique sys-tem with regard to their evolution through landslidedevelopment. Magma-jacking in the rift zonesŽ .Swanson et al., 1976 and consequent seismic trig-

Ž .gering of larger earthquakes Bryan, 1992 causesslippage. This slippage may take place along the

Ž .volcanorcrust interface Lipman et al., 1985 , basedŽon earthquake foci Klein et al., 1987; Thurber and

. ŽGripp, 1988 and seismic refraction experiments Hill.and Zucca, 1987 . This gliding may occur either on a

Žbasal layer of either pelagic sediments Nakamura,. Ž1982 or a dunite cumulate body Clague and Den-

.linger, 1994 . However, the recent finding that asedimentary wedge underlies a large portion of off-

Ž .shore Kilauea Naka et al., 1998 suggests that inter-mediate depth failure along an indurated sedimentlayer may also be possible. Both deep and shallowslip panes have been proposed for the south flank

Ž .based on seismicity e.g., Lipman et al., 1985; deepŽ .and leveling data Arnadottir et al., 1991; shallow .

In this paper, we have shown how the resultantslump mass exhibits the surficial manifestations of

Ž .the rotational slumping model of Varnes 1978 ,Žincluding some internal aspects of large deltas De-

.nlinger and Okubo, 1995 along with evidence ofthrusting within the main slump body and at the toeŽvan Bemmelen, 1949; Varnes, 1958; Borgia et al.,1990; Borgia and Treves, 1992; Denlinger and

.Okubo, 1995 .A comprehensive theory we favor for explaining

the features evident in this complex volcanic systemis one with volcano spreading as the principal defor-mational process. The volcano spreading process was

Žrecently applied to Hawaiian volcanoes Borgia et.al., 1990; Borgia and Treves, 1992; Borgia, 1994 ,

and leaves room for smaller, intermediate scale flankfailures to occur, as described in this paper. Thissmaller class of failures may not be as wide inextent, and may sole closer to the surface than thelarger-scale process of volcano spreading describedin the Borgia models. Preliminary results from the

Žrecent KaireirKaiko surveys Naka et al., 1998;Shipboard Scientific Party, 1998; Lipman et al., sub-


.mitted show that we still have much to learn aboutthe structure and evolution of Hawaiian volcanoes,giant landslides, and oceanic islands in general.

5. Hazard assessment for the Kilauea south flank

What can we deduce about the triggering of slopeinstabilities and subsequent driving forces of flankfailures in the Hawaiian Islands? Maximum averageslopes of the islands are 138 onshore and 178 off-

Žshore Mark and Moore, 1987; Moore and Chad-.wick, 1995 which are stable according to the

geotechnical literature, even when considering poreŽpressure effects Terzaghi and Peck, 1967; Zaruba

. Ž .and Mencl, 1982; Rice, 1985 . Iverson 1995 per-formed an extensive rigid-wedge limit equilibriumanalysis specifically focused on magma-injection andgroundwater forces as potential triggering mecha-nisms of the giant Hawaiian landslides. After dis-cussing the effects of existing slopes and gravity,magma-pressure, and various groundwater-seepageforces, he was unable to conclude that any of theseeffects alone could produce such major events exceptunder restrictive circumstances that do not match the

Ž .available observations Iverson, 1995 .Gravitationally induced slides are marked by pre-

cursors that may precede failure by months or yearsŽ .Rice, 1985 . Slope failure of a cohesive material isusually preceded by the formation of tension cracks

Žbehind the upper edge of the slope Terzaghi and.Peck, 1967 , analogous in the present study to the

ŽKoae fault zone Moore and Krivoy, 1964; Denlinger.and Okubo, 1995 . Theories have been developed

based on the assumption that two phases contributeto this type of sliding. The first is a visco-plasticstage of creep, which we infer is the present state ofthe south flank. This is followed by a second phase

Ž .of abrupt, brittle fracture Jaeger, 1979 which canŽ . Ž .produce catastrophic slip. Rice 1985 p. 307 pro-

vides numerous examples of noted landslides, includ-ing the Vaiont slide, which exhibited movementandror fractures prior to a conclusive catastrophicslip.

Research on seismically induced slips has shownthat a considerable duration of ground motion is

Žrequired to initiate a landslide e.g., Seed and Good-man, 1963; Ellis and Hartman, 1966; Seed and Wil-

.son, 1966 , which typically shows little precursoryŽ .evidence of impending failure Rice, 1985 . Ground

shaking needs to occur on the order of minutes toactivate such a landslide, and slope movement hasbeen found to cease immediately after earthquake

Ž .motion stops Rice, 1985 . For example, the strongshaking of the 1975 Kalapana earthquake main shockwas reported to have lasted 50 s with an average

Žrupture velocity of 0.8 kmrs Harvey and Wyss,.1986 . Questions arise as to how and why such large

earthquakes occur and do they trigger or result fromthe landslide. Precursory shocks occurred prior to the1975 Kalapana main event, so it is obvious thatearthquakes played a part in the initialization offlank movement. However, the question is: was themain shock predestined at 7.2, or was the mainshock actually smaller but amplified by the effects ofthe sliding, rotating, or translating south flank? BryanŽ .1992 proposes that small earthquakes, releasing thelocal normal stress concentration regime from com-plex rift zone processes such as magmatic injection,in turn trigger the large-scale energy release at thebase of the volcanic edifice manifested by low-anglesliding events. This requires the local stress regimeto be superposed upon the regional stress field result-ing from both volcanic loading of Hawaii Island andits consequent lithospheric flexure.

Ž . Ž .Wood 1914 and Wyss et al. 1992 reported thathundreds of foreshocks and aftershocks over severaldays were associated with the 1868 earthquake, thepeak of which occurred in several large shocks in-stead of one sustained or extended event. Thoughflank motion downward and seaward has occurred

Ž .during these major seismic events cf., 1868, 1975 ,perhaps this pattern of more gradual stress releasehas kept the entire Hilina slump block intact to date.Since there are two types of landslides on the Hawai-

Ž .ian Ridge i.e., slumps and debris avalanches , itmight be possible that they are triggered by differentmeans. The continuous creep of the Hilina slumpmay be largely driven by the volcano spreadingprocess consisting of the forces of magma-intrusion

Ž .and gravity proposed by Borgia 1994 , with precur-sory clues common to gravitationally driven failuresas previously discussed. Hawaiian debris avalanches,

Ž .such as the Nuuanu slide Moore et al., 1989 , mayinstead result from seismic triggers since theirwidespread and far-reaching deposits are the prime


Ž .characteristics which Solonenko 1977 found to dis-tinguish seismically induced events. Another possi-bility is that of an explosion within the volcanoinitiating flank collapse and a subsequent debrisavalanche. Such scenarios were proposed by RiceŽ .1985 for the Mount St. Helens north slope failure

Ž .and McMurtry et al. 1999-this volume for the Alika2 slide.

The most worrisome potential with respect to thefuture of the Hilina slump may lie with a combina-tion of the two triggering processes of magma-jack-ing and seismicity. First, magma-jacking would causegradual development of the breakaway lateral bound-aries by creep with interspersed large flexural-relatedlithospheric earthquakes giving boundary develop-ment an added push. Then, once the boundaries arefully developed and easily slipped, less stress isbuilt-up by magma-jacking. The flank would now bepoised for imminent catastrophic failure in the formof major slumping events or a debris avalanchecoincident with the occurrence of a lithosphericearthquake of large magnitude and extended dura-tion. Presently, most study and attention of southflank movement is focused on the southwest bound-ary because of its high displacement rate. We shouldalso focus attention and monitoring efforts on thesub-aerial, and possibly submarine, northeast bound-ary to best determine the potential hazard. Owen et

Ž .al. 1995 comment that no displacement patternsindicative of a tear fault were evident there during

Ž .their 3-year study Fig. 4C . They suggest the north-east part of the flank may currently only moveseismically over greater time periods. Additionally,instrumenting the submarine southwest boundary aswell as the bench would be prudent.

The phrase ‘‘catastrophic failure’’ in this contextoften conjures images of half the volcano calving offand sliding into the sea in one instantaneous prodi-gious event, similar to the scenario proposed fordebris avalanches such as the Nuuanu slide on OahuŽ .Moore et al., 1989, 1994a,b; Normark et al., 1993 .An event of this magnitude would indeed be devas-tating to human life and property as well as to theisland itself, though such an extreme level of dias-trophism is not required to wreak havoc on thepopulation. The large south flank earthquakes in the

Ž .historic past 1868, M;7.9; 1975, Ms7.2 haveresulted in deaths, extensive property damage, and

meters of sudden flank movement along the Hilinafault system. Just a two-fold increase in the amountof seaward and downward slope displacement froman earthquake of larger magnitude and longer dura-tion than the 1975 event would truly be disastrous tolife and property on Hawaii island, the rest of thearchipelago, and possibly the Pacific Rim.

In the island’s frame of reference, however, tensof meters of displacement to an edifice that standsnearly 10-km high and is almost 200 km in diameteris trivial and merely part of the evolution of thevolcanic complex over geologic time, on the order ofhundreds of thousands of years. Yet, a paroxysmaldebris slide removing up to 25% or more of anindividual volcano would be catastrophic to thestructure itself, drastically changing its plumbingsystem, exposing its inner core, and perhaps drivingthe volcano beyond the peak of shield-building intothe waning next stages. Such an event, possibly animmense phreato-magmatic eruption, and the result-ing natural disasters would be extensive, requiring an

Žearthquake estimated at magnitude 8.8 Klein and.Okubo, 1993 , and may better be described as

‘‘cataclysmic’’ in human terms. Though it is yetunclear if slumps such as those on the south flankdegrade into cataclysmic debris avalanches, as pro-

Žposed by some researchers Moore et al., 1989;.Normark et al., 1993 , it is not required in order to

be catastrophic on a human scale. The episodiclarge-scale flank readjustments experienced throughhistoric time and prehistorically evident in the pres-ence of the grandiose 500-m-high Hilina Pali, whichmay possibly grow only by such abrupt displace-ments, demonstrates that volcanologically minorevents can indeed be humanly calamitous.

6. Conclusions

Regional oversteepened slopes are not common toHawaiian volcanoes in the growth phase. Slumpspreferentially form in secondary volcanoes whichgrow on the flank of a primary edifice. The recentfinding of thickened indurated sedimentary wedges

Ž .beneath a newer volcano Kilauea may play a keyrole in eventual flank failure, either catastrophic orgradual. Triggering mechanisms, in addition to gravi-tational forces, must act to initiate major slope fail-


ures. Buttressing pins are set in the form of ancientand active seamounts, affecting rift zone and slopedevelopment. New seamounts, such as Loihi, mayimmediately serve the competing roles of a budding,seismically active volcano and a slope stabilizingabutment.

Based on results from the new high-resolutionoffshore mapping, the estimated area presently af-fected by south flank slumping is roughly 2100 km2,slightly larger than the sub-aerial portion of Maui. Ifone assumes slippage along listric faults down to thebase of the volcanic pile at ;9-km deep, an approx-imate slump volume of 10,000–12,000 km3 is delin-eated, which translates to about 10% of the entireHawaii island mass. The Hilina fault system does notcontinue offshore as a series of normal faults down-thrown seaward; instead, the submarine flank com-prises four distinct zones, broadly in line with the

Ž .Varnes 1978 rotational slump model. Large elon-gate ridges at the base of Kilauea volcano are inter-preted as detached slide blocks from the outer scarp,both consisting of pre-Kilauea sedimentary rock.These sandstone–mudstones, and the outer scarpmorphology, were likely formed by compressiveforces generated by Kilauea overthrusting the sedi-ments while growing seaward. The blocks wereprobably abruptly emplaced during major seismicepisodes or by creep accelerated over the rate ofmovement of the larger slumping south flank mass.

Once the northeast lateral boundary of the Hilinaslump develops to the same extent as the possiblyfully evolved southwest boundary, then an increasedlikelihood of potentially catastrophic failure may oc-cur. These events could be triggered by Hawaiiisland earthquakes of large magnitude and extendedduration.

Acknowledgements

This study was supported by NSF grant OCE-9116535 to ANS and AM and additional NOAAsupport to JRS through AM’s Sea Grant andNURPrHURL grants. Special thanks to the captains,crews, and technicians who carried out the variousfield programs and the USGS for cooperation in theHAWAII MR1 survey under an MOU with SOEST.A contract with the Far East Division of the former

Soviet Academy of Sciences allowed for collectionof the seismic reflection data. The Japan Marine

Ž .Science and Technology Center JAMSTEC pro-vided the resources to carry out the RrV Kairei andROV Kaiko surveys and the Shipboard ScientificParty was composed of scientists from Japan,SOEST, and the USGS. Some of the SEA BEAMdata were formatted for our use by A. Bobbitt ofPMEL and W.W. Chadwick of OSU. HMRG andHURL provided computer resources. Appreciationgoes to A. Rice for calling to our attention hisstudies of the Mount St. Helens landslide and to S.Martel for pointing out additional geotechnical litera-ture. J. Kauahikaua of HVO and J. Foster of SOESTprovided the digital database of sub-aerial geologicstructures for our figures. T. Duennebier was instru-mental in the initial preparation of the topographicand bathymetric grid. Thorough reviews by C.M.Riley and another referee greatly improved themanuscript. G.M. McMurtry also offered construc-tive advice and D. Henderson provided editorialassistance. SOEST contribution number 4849.

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Submarine geology of the Hilina slump and morpho-structural evolution of Kilauea volcano, Hawaii