Crustal Structure of the Tuamotu Plateau, 15&Deg

JOURNAL OF GEOPHYSICAL RESEARCH, VOL. tOO, NO. B5, PAGES 8097-8114, MAY t0, 1995

Crustal structure of the Tuamotu Plateau, 15øS, and implications for its origin

Garrett Ito MIT/WHOI JointProgram in Oceanography,Woods Hole OceanographicInstitution, Woods Hole, Massachusetts

Marcia McNutt Departmentof Earth,Atmospheric, and Planetary Sciences, Massachusetts Institute of Technology,Cambridge

Richard L. Gibson Earth ResourcesLaboratory, Department of Earth, Atmospheric,and Planetary Sciences MassachusettsInstitute of Technology,Cambridge

Abstract. We investigatethe sedimentaryand volcanicstructure of the TuamotuPlateau with multichannelseismic, seismic refraction, and gravity data along a shiptrack crossingthe plateau near 15øS. The volcanicbasement of the centralportion of the plateauis cappedwith a 1 to 2- km-thicksediment layer composedof two compositionalsequences. The uppermostsequence, with semblance-derivedP wave velocities of 1.6-1.9 km/s and thicknessesof 0.2-0.9 km, is composedof pelagic sediments.The underlyingsequence, with velocities2.5-3.5 km/s and thicknessesof 0.5-1.5 km, is composedof limestoneand volcaniclasticsediments. Sonobuoy refractiondata showthe upper 1 km of the volcanicbasement to have velocities4.5-5.5 km/s. The gravity dataindicate that the platformis compensatedby an elasticlithosphere with effective thickness 5+5 km and that the volcanic thickness is 9-10 km thicker than normal oceaniccrust with a volumeof 2.0-2.6x106km 3. Theinferred eruption rates of 0.1-0.13km3/yr are comparableto thoseof the Hawaiian andMarquesas island chains but substantiallyless than thoseof many oceanicplateaus. Radiometric and paleontological ages for the plateauand geomagneticdates of the surroundingseafloor indicate that the northwesternportion of the plateauformed -600 km off the axis of the paleo-Pacific-Farallonspreading center, on lithosphereof age ~ 10-20 Ma. Linear volcanicridges and scarpsbounding deep sediment-filled basins,however, are similarto featuresof oceanicplateaus which formedat or near accretionary plateboundaries. We suggestthat these volcanic ridges and the grossplateau like morphology were formedby magmathat was channelledalong the lithosphericdiscontinuities left behindby a southwardpropagating rift segmentof the nearbyspreading center. We attributethe formation of the northwesternportion of the TuamotuPlateau to the passageof two hotspotsduring times 50-30 Ma as they migratedbeneath the Pacific platebut remainedwest of the Tuamotu propagator.

Introduction areless than magnetic ages of theunderlying lithosphere [Jarrard and Clague, 1977; Henderson,1985], mostisland chains were Hotspotsin the mantle, are thoughtto be the sourcesof many formedmidplate, far from oceanicspreading centers. As a result, crustal anomaliesin the world's ocean basins. Oceanic hotspot most island chainsare compensatedby mechanicallycompetent features can be separatedinto two classes: island chains and lithospherewith effectiveelastic thicknesses (typically 10-25 km) oceanic plateaus. The morphological characteristics that thatthicken with age,consistent with a conductivecooling model distinguishthese two classesmay reflect differences in the for the lithosphere[Watts, 1978; Watts et al., 1980]. tectonic environments at which they formed and/or the mantle Oceanicplateaus, on theother hand, are defined by theirbroad sourceswhich producedthese melt anomalies. Ocean island elevated surfaces which, in many cases (e.g., Ontong Java, chains are composed of discrete volcanic edifices with ShatskyRise, and Manihiki), are nonlinear in planview andlack geographic age distributions reflecting the motion of the the geographicage distributions so typical among island chains. lithospherewith respectto the hotspotreference frame [Duncan Most oceanicplateaus are thoughtto have originatedfrom and McDougall, 1976; Clague and Jarrard, 1973; Morgan, hotspotssited near or at oceanicspreading centers or rifted 1972]. As demonstratedby isotopicand paleontologicalages that continentalmargins because (1) edificeages are close to thoseof the surroundingseafloor [Detrick et al., 1977]; (2) subsidence Copyright1995 by the AmericanGeophysical Union. histories of oceanic plateaus match closely those of the surroundingseafloor [Detrick et al., 1977]; and (3) oceanic Papernumber 95JB00071. plateausare compensated by Airy isostasyindicating they loaded 0148-0227/95/95JB-000710505.00 thin elastic lithospheresfound near ocean spreadingcenters

8O97 8098 ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

[Detrick and Watts, 1979; Kogan, 1979; Sandwell and easternportion of TuamotuPlateau may have beenemplaced near MacKenzie, 1989]. In addition, plateaus such as Manihiki the paleo-Pacific-Farallonspreading center, as suggestedby the [Winterer et al., 1974] and Kerguelen [Munschy and Schlich, underlying weak lithospherewith effective elasticthickness of 2- 1987] show faulted structureswhich may have resulted from 6 km [Okal and Cazenave, 1985]. extensionalprocesses at an oceanicspreading center. The broad The nearly parallel alignment of the Line Islands-Tuamotu continuousextent of oceanicplateaus may result from efficient chain with the Hawaiian-Emperorchain led Morgan [1972] to penetrationof heat and magmathrough the weak lithosphereat hypothesize that the two were formed concurrentlyby two oceanicspreading centers [Vogt, 1974; Okal and Batiza, 1987]. stationary hotspots. He suggestedthat the northwesternmost Volcanic volumesof oceanicplateaus, representing 5-25% of the portionof Tuamotuplatform marked the changein plate motion world's total volcanicproduction for the past 150 m.y. [Larson, recordedby the bend in the Hawaiian-Emperorchain datedat 42- 1991], are typically much larger than the volumes of island 43 Ma [Clague and Jarrard, 1973]. Subsequentpaleontologic chainsand may be the productsof massiveflood basalteruptions and radiometric age constraints, however, indicate that the generatedat the onsetof mantleplume activity [Morgan, 1972, northernmost Tuamotu islands formed as late as 47-55 Ma 1981; Richards et al., 1989]. [Martini, 1976; Jacksonand Schlanger,1976; Schlangeret al., The Tuamotu Plateau, in the south central Pacific, is a volcanic 1984], thuspreceding the Hawaiian-Emperorbend by as muchas feature with characteristics of both island chains and oceanic 9 m.y. Several hotspotsare suggestedto have formed the Line plateaus(Figure 1). From a regionalperspective, the platform Islandschain, based on the geochronologicstudies of Schlanger appearsplateau like, with a broadelevated surface of width80- et al. [1984] and Winterer [1976], while two hotspots are 200 km. A closer look, however, reveals discrete volcanic hypothesizedto have formed the Tuamotu Islands from tectonic pinnacles,the largestof which are now reef-coveredatolls. The studiesof Okal and Cazenave[1985]. Thus, the origin of the platformis a linear featuretrending from northwestto southeast platform may be more complex than the single hotspot similar to Pacific island chains such as the Hawaiian chain, the mechanismoriginally suggestedby Morgan [ 1972]. Cook-Austral chain, and the Society Islands. However, the In this paper we investigate the sedimentary and volcanic

b) -10ø / •.• '--;•--,. '- Marquesas.• ,, •- • • -• !"•'- •C•'•'•.•:...... • Tnamotn '•, •'• • •- !

I• •s•:•:•:•.'•.-3 Islands • •,..•.'..,..... •" '.,'½•?, ;•.... • '•:.,> '• ½•' •.•'"-•.,-, • ', , x H

•,•., • ....• •.i'• • x•' x_,•--•, '--,, ' x....

_20• • -,.%-.,. •h --W.-•.>.•,.-., ......

[•:' • Austral ".• • - •e/" •-'=•"..... • c• • ..... l

I / 205 ø 210 ø 215 ø 220 ø 225 ø Longitude Figure 1. (a) Regionalmap surroundingthe TuamotuPlateau contouring the 3-km isobath. The rectangle outlinesthe bathymetrymap below. (b) Bathymetrymap of the Tuamotuplatform with 1-km contourinterval. The solidline marksthe surveyline whichcrosses the northernportion of the platformin the regionmarked by the rectangle. ITO ET AL.- CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8099

_14 ø

_15 ø

211 ø 212 ø 213 ø 214 ø Longitude

-5000 -4000 -3000 -2000 - 1000 -0 Depth(m) Figure2. Bathymetrymap of thesurvey region (rectangle in Figure lb) witha contourinterval of 500m. The locationsof thesonobuoy deployments are marked by X' s alongthe survey track shown as solid white. Sections A, B, andC of thesurvey line are sediment basins revealed in thereflection profile. Dashed lines mark lineations of volcanicridges and scarps inferred from the volcanic morphology observed in thereflection profile.

structure of the northwestern end of the Tuamotu Plateau in order Gravity data were collectedwith a KSS-30 gravimeter,while to better understand the tectonic and magmatic processes bathymetrywas determined from the centerbeam readings of the controllingits formation. We imagethe shallowcrustal structure Hydrosweepmultibeam system. The bathymetrysurrounding our (upper-2 km) with multichannelseismic (MCS) reflectiondata shiptrack are digitized points from bathymetrymaps produced at obtainedalong a ship crossingof the northernportion of the the Institut Francois de Rechere pour l'Exploitation de la Mer platform. The velocitystructure used for stackingthe reflection (IFREMER) and are from the Digital BathymetricData Base 5 dataand for interpretationof sedimentarycomposition is derived (DBDB5) data set (Figure 2). Also shownin Figure 2 is the from semblanceanalyses of the MCS data and from sonobuoy locationof Deep Sea Drilling Project (DSDP) site 318, which refraction data. To constrainthe deep crustal structureand penetrated745 m into the sedimentarylayer of thenorthern flank strtng•hof the compensatinglithosphere, we usegravity and of the platform. Correlationof the drill core findingswith our bathyrnetrydata. While the volume and flux of Tuamotu velocityresults will be discussedbelow. volcanismmay havebeen controlled by the mantlesource which wasof Comparablestrength to the Hawaiianand Marquesas Seismic Reflection Data Analysis h0tspots,the volcanicmorphology of the Tuamotuplatform may havebeen formed by structuraldiscontinuities in the lithosphere, Prestack Processing engravedby a propagatingsegment of thenearby Pacific-Farallon Before stackingthe MCS data, we performedthree processing spreadingcenter. steps using the software package ProMax (ProMax 40.22, AdvanceGeophysical Corporation, copyright 1989-1992). In the Data Collection first step,we applieda band-passfilter (5-62 Hz) to eachchannel To reveal subsurfacestratigraphy of the sedimentaryand recordto eliminatelow-frequency streamer noise and undesired volcaniclayers, we obtainedmultichannel reflection data along high frequencies.The secondstep was to filter the shotgathers in the surveyline shownin Figures1 and 2. We usedthe MCS frequency-wavenumber(f-k) spaceto eliminate energy which systemon boardthe R/V MauriceEwing, which consisted of a coincidedwith apparentvelocities less than the water velocity 20-airgun(8385 in 3) arraywhich shot to a 148-channelstreamer (1.5 km/s), thereby reducing coherent noise seen as steeply with maximumchannel offset of-3.7 km. Shot spacingwas -50 dippinglines in unfilteredstacked data. Suchcoherent noise may m,yielding 37-fold, common-midpoint shotgathers; the sampling have arisenfrom wavestraveling from sourceto receiverthrough interval was 4 ms. The line was shotwhile steamingnortheast to the streamercable or guidedin the watercolumn, or by scattering southwest.To constraindeep crustalvelocities, we launched from out-of-planepoint sources [Larner et al., 1983]. three expendablesonobuoys (buoys 1-3, Figure 2) which The final step in the prestackprocessing was predictive recordedrefractions of the air-gunsignal out to offsetsof-20 deconvolution. For each channelrecord, we specifiedtwo time km. gatesfrom whichwe designeddeconvolution filters. We setthe 8100 ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

o o o o o o o o. o. o. o. o. o. o. ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8101 length of the upper gate to be between0.5 and 1.5 s, so as to RegionsA, B, andC differfrom the rest of theplateau in that include the seafloor arrival and shallow sedimentlayers, and the reflectionsare nearlyhorizontal, smoother, and more continuous length of the lower gate to be between 1.0 and 2.0 s, to include (Figure4). Layer 1 is boundby the seafloor(reflector a) and arrivalsbelow the first gate but abovethe seafloormultiple. We reflectorb, while layer2 is boundby reflectorsb andc. In region appliedthe solutionof theupper gate to datawithin andabove the A, layers 1 and 2 are approximately the same thicknesswith uppergate, and the solutionof the lower gate within and below combinedthickness -0.7 s. In region B, layer 1 is 0.2-0.4 s thick the lower gate. This two-gate procedurewas chosento account and overlies layer 2 which thickenseastward from 0.2 s to 0.9 s. for changesin frequencycharacter with increasingtravel time, In region C, the two layers extend as deep as -2.3 s below the presumably due to both changesin sediment lithology and seafloor. increasedattenuation of the higherfrequencies with depth. Region A and the eastside of regionC are boundby scarpsof A preliminary stackingwithout deconvolutionproduced a pinnaclelike structures,the peaksof which are cappedby very seriesof reverberationsbeneath many of the primary reflectors little sediment and reveal no reflectors below their opaque which, in the Fourier transformdomain, showedas peaksat 13, surfaces.The seismicrecords within thesethree regionsresemble 18, 25, 32, and 36 Hz. The direct water arrival signal had a those of lagoonal sequencesover the Mid-Pacific Mountains similar amplitudespectrum, suggesting that many of the signals observedby Wintereret al. [ 1993], suggestingthese stratigraphic in the stackedsection were artifactsof incompletecancellation of sectionsare basins of carbonatesediments filled by continuous the bubble-pulse signals from the airgun source array. sedimentinflux from the surroundingvolcanic slopes. Reverberationsobserved deeper in the stackedsection might have also been generatedby multiple reflections within subsurface Velocity Analysis sedimentary layers (i.e., peg-leg multiples). Our predictive Accurate constraints on crustal elastic wave velocities are deconvolutionmethod proved effective in suppressingthese important for high-quality stacking of the MCS data and for undesiredsignals. interpretation of sedimentary and crustal composition. Semblancecalculations were madeevery 10 CMP' s, then two to Results five adjacentcalculations were stackedto enhancethe signalto After applying deconvolution, the data were sorted to noiseratios. We usedthese stacked semblance results to pick common-midpoint gathers (CMP's), corrected for normal root-mean-square(rms) velocities. moveoutusing semblance-derivedvelocities, and then stacked. To constraindeeper crustal velocities, we usedwide-angle The stacked time section is shown in Figure 3a, while the refractionsrecorded by two sonobuoysover the central portion of reflectionswe identify are shownin Figure 3b. Between regions the plateauin additionto the semblanceanalyses. Sonobuoy 1, B and C (211.6ø-212.9øE which we will refer to as the central launched near 211.9øE, recorded refractions out to an offset of portion) the plateauis drapedby a layer of thickness0.2-0.6 s -22.0 km as the shipsteamed west. Sonobuoy2, deployed-30 (two-way travel time) characterized by rippled, but nearly km east of sonobuoy 1, recordedrefractions out to an offset of continuousreflections (layer 1). The lower boundaryof this layer -17.5 km. To determinecrustal velocities, we performedforward is the horizon immediately below the seafloormarked in Figure modeling by one-dimensional(l-D) raytracing using the 3b. Below this horizon is layer 2, characterizedby slightly less inter•.ctiveraytracing feature of JDSeis(JDSeis, by JohnDiebold, continuousreflectors and more high-frequencynoise between Lamont Doherty Earth Observatory,Palisades, New York, 1992) reflectorsrelative to layer 1. The lower boundary of layer 2 is andtwo-dimensional (2-D) raytracingusing the softwarepackage definedby the deepestreflections in regionsmarked A, B, andC, MacRay [Luetgert, 1992]. Before the raytracingwas done, but is uncertainwithin the centralportion of the plateau. In the however, sonobuoyrecords were correctedfor topographic centralportion of the plateauwe observereflections (R1, Figure effects using the method describedby Officer and Wuenschel 3b) thatnearly parallel layer 1, 0.6-0.9 s below the seafloor.The [1951]. deepestreflections we identify are a numberof discontinuous Figure 5 showsthe sonobuoyrecords and raytracingcurves reflections(R2), 1.6-2.5 s below the plateausurface. The crustal with their correspondingvelocity profiles. The directwater wave velocities shown in Figure 3b and the depth sectionshown in is seenas the linear arrival in the sonobuoyrecords. Arrivals a, Figure 3c are discussedlater. in the sonobuoyrecords, are reflections from the base of the

Figure 3. (opposite)(a) The stackedtime sectionof the TuamotuPlateau shown from west to east which includes25,780 CMP tracesspanning -360 km. For plotting,we stackedevery four tracesand applieda cubic polynomialgain function. The coursechanges shown in Figure2 aremarked by thebold vertical lines while the regionssampled by sonobuoysare marked by the shortvertical lines. (b) Time sectionshowing major reflections and semblance-derived interval velocities. The shallowest subsurface reflections correlate with the base of the lowest velocity layer (layer 1, white). The seafloormultiple is outlinedin white. The centralportion of the plateau (211.6ø-212.9øE)is between regionsB and C. (c) A depth sectionplot obtainedfrom semblance velocitiesshowing reflections and sedimentcompositions inferred from findingsat DSDP site 318. The upper pelagiclayer is white, while the lower limestone/volcaniclasticlayer is dotted. The hachuredregion is the 4.65 km/s velocity layer which most likely enclosesthe volcanic basementinterface in the central portion of the plateau.The shadedregion marks the volcaniccrust. ReflectionsR1 maybe from the volcanicbasement, while reflectionsR2 may mark a compositionalvolcanic boundary. Atolls, seamounts,and DSDP site 318 are marked at locationsperpendicular to the shiptrack. 81o2 ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

a) 2.0

3.0

4.0

5.0

211.0ø Longitude 211.5ø I 50 km

b) 3.0

4.0

5.0

6.0

213.0 ø 213.5 ø Longitude Figure 4. Reflectiontime sectionsmarked along major reflectorsdesignated by lettersfor (a) regionsA and B and (b) regionC. Reflectiona is the seafloor,reflection b marksthe baseof layer 1 (the low velocitysediments), and reflectionc marks the volcanicbasement interface, also the lower boundof sedimentlayer 2. Bounding scarpsand volcanic mounds are linearfeatures seen in the surveymap of Figure2.

watercolumn (velocity 1.5 km/s). Arrivalsb, immediatelybelow derived interval velocitiesat the midpoint depthsof velocity the seafloor arrival, are from the base of a low-velocity layer layers for semblance calculations within the range of each (1.6-2.0km/s, layer 1) with approximatethicknesses of 0.35 km sonobuoy record. The raytracing methods and semblance below sonobuoy1 and0.48 km below sonobuoy2. analysesproduce consistent results for the first -1.5 km (-1.0 s) In the sonobuoy1 record,deep refractions cross the seafloor below the seafloor; however, semblancevelocities are slightly arrival near an offset of 3.6 km. These refractions are from a higherthan raytracingvelocities at depthsgreater than -1.5 km high-velocityregion in which we distinguishtwo layers: (1) beneaththe seafloor. The consistencybetween semblance results layer2, of thickness-1.4 km andvelocity 3.7-4.0 km/s, evident and sonobuoyresults in the upper 1 s strengthensour confidence by first arrivalrefractions recorded out to an offsetof-10 km; in semblance-derivedvelocities above this depth. and (2) layer 3, with thickness>1.8 km and velocities4.3-5.0 We show semblance-derivedvelocity layers for the whole km/s evidentby first arrivalsrecorded at offsets>10 km. The platformin Figure 3b. Layer 1 identifiedin the stackedsection combinedthickness of the two layersis at least -3.5 km. The correlateswell with the interval of velocity less than 1.9 km/s high-velocity region sampled by sonobuoy2 has a more shown as the white layer immediately beneath the seafloor. continuousvelocity profile which we identifyas a singlevelocity Layer 2 is composedof two velocity layers: the first with layer (layer 2). Refractedenergy from this layer first intersects averagevelocity 2.5 km/s and thickness0.1-0.4 s, andthe second the seafloor arrival, 2 km further in offset than in the sonobuoy1 with average velocity 3.5 km/s and thickness0.1-0.5 s. An record,indicating a moregradual increase in velocitybelow layer interval with average velocity of 4.65 km/s occurs 1.0-1.7 s 1. Layer 2 velocitiesincrease to -5.4 krn/sover a thicknessof below the seafloor. This interval is the last we chose to shade -3.4 km. The transitionbetween layer 2 and layer 3 obtainedby because its velocity is that at which the semblance-derived sonobuoy1 may be the sediment-volcanicbasement transition. velocities began to deviate from sonobuoy-derivedvelocities. This transition,however, is less evident in the sonobuoy2 record The velocity of 4.65 km/s is alsothe velocitythat Talandier and becausethe sedimentsoverlying the volcanicbasement in this Okal [1987] attributeto the extrusivebasaltic layer from their region have greatervelocities, approaching those of basement, regional refraction study. Depth uncertaintiesfor the intervals possiblydue to highervolcaniclastic content. range from 0.1 km for the 1.9-km/s velocity contour,to 0.5 km For comparison,we also plot in Figure 5 the semblance- for the 4.65-km/s contour. ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8103

Offset (km) Velocity (km/s) 5 10 15 20 25 1 2 3 4 5 6 a) •.La_y_e r_1_ _ \ ':"'•'.o'':.? .•..,,,. Layer- 2 ,, X '.;.

\\••;•[.. .':'... Layer3•.' '..-'

ß o _ \\

5 10 15 20 25 I I I I 1 2 3 4 5 6 b) 2

2 - , L.,a.ye_r. 1 -"- -:E •'-',:.•,-..'.. •' 3 ' x.;:x'

.o...ß. ß 5 ß o 4 - XX•ß . ß b 6 Layer 2 .'

\\ 7 \x 6 Figure 5. Sonobuoyrecords outlined with the 1-D raytracingcurves, and corresponding velocity profiles for (a) sonobuoy1 and(b) sonobuoy2. The solidvelocity profiles are the 1-D raytracingsolutions which are consistent with the shallowestand deepest velocity nodes of the 2-D raytracingsolutions connected by the dashedlines. The dotsmark intervalvelocities at the midpointdepths of eachvelocity layer for semblancecalculations within therange of the sonobuoyrecords. Semblance velocities begin to deviatefrom raytracing solutions at -1.5 km below the seafloorand neara velocityof -4.5 krn/s. Arrivals a are the seafloorreflections and arrivalsb, just below arrivals 1, are the baseof the low-velocity sedimentlayers.

Sediment Composition and Volcanic Morphology plateauas the dottedlayer in Figure 3c with thickness-0.5 km. The compositionaldifference between layer 1 and2 may explain From these constraints on crustal velocities and from the previouslynoted differencein reflectioncharacter between stratigraphy imaged by the stacked MCS data, we infer the two layers. composition of sediments and structure of the underlying Within the layer of interval velocity 4.65 km/s lies the volcanics. We baseour interpretationsof sedimentcomposition sediment-volcanicbasement transition (with the exception of on the findingsof DSDP site 318 which penetrated745 m below region C to be discussedbelow) shownas the hachuredregion in the seafloor, -25 km southeastof our survey line near 213øE Figure 3c. Through most the central portion of the plateau, the [Schlangeret al., 1976] (marked in Figure 3c and in the map in volcanic basement interface is not evident in the reflection Figure 2). The compositionsof the drill core sections are section presumably because of the small (or nonexistent) outlined in Table 1. The velocities of layer 1 correspondwell impedance contrast between the deep sedimentaryrocks and with those for units 1 and 2 of the drill core; therefore, we surfaceextrusives. ReflectionsR1 (Figure 3), however,might be interpret this upper layer to be pelagic ooze becoming more from this interface. The inferred volcanic basement between lithified with depth to form chalk. This is shown as the white 211.75øE and 213.0øE graduallydeepens from west to east with layer in Figure 3c draping the central portion of the plateau very little apparent relief. The total sedimentthickness in this (211.6ø-212.9øE) with thickness 0.2-0.9 km. The velocities of region is 1-2 km and suggestsan averagesedimentation rate of layer 2 (2.5-3.5 krn/s)correlate well with thosefor the lower two 20-40 rn/m.y., assuminga basementage of 50 m.y. The bulk of units of the drill core; therefore, we believe this layer is this layer may have beenbuilt as reefs when the centralportion of composedof chalk near the top, and limestonewith increasing the plateau was near sea level. Then, as the platform subsided, concentrations of volcaniclastic sediments near the bottom. This the reefs drowned and were blanketedby pelagic sedimentsand compositional layer is shown over the central portion of the sedimentseroded from higher-standingedifices. 8104 ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

Table 1. RecoveryFrom DSDP Site 318

Unit DepthRange, m Composition CompressionalSound Age Velocity, km/s

1 0.0-265.5 foram-nannofossilooze grading 1.6-2.0 Pleistocene to soft chalk, volcanic sand

2 265.5-416.0 brecciaand conglomerate basalt 1.7-2.0 lowerOligocene- clasts,foram-nannofossil chalk, upperPliocene volcanic sand, and chert

3 416.0-530.0 foram-nannofossillimestone 2.0-3.5 upper Eocene rich in reefal forams; corals, volcanic siltstone and sandstone

4 530.0-745.0 nannofossillimestone, siliceous 2.0-4.4 lower Eocene- volcanic sandstoneand siltstone middle Eocene

Geologicaldata are from Schlanger etal. [1976],ages are from Martini [1976], and seismic velocities are from Boyce [1976]

Nearstratigraphic regions A, B, andC, however,the volcanics consistentwith our requirementthat velocitiesexceed the 5.5- formsteep-sided mounds with little sedimentcovering over the krn/s maximum observedin the sonobuoyrecords to producean peaks.The sediment basin in regionA is boundedon the west by impedancecontrast sufficient to yield reflections.We further a volcanicridge forming the western margin of theplateau and to investigatethe deep crustalstructure from analysisof gravity theeast by theslope of Mataivaatoll (see Figures 2 and3). The data. volcanic basementforms a valley which has accumulateda sedimentlayer of thickness-0.7 km. Again,assuming an age of CompensationMechanisms and Deep Crustal 50 m.y.for theunderlying volcanics, the average sedimentation Structure ratefor basinA is -14 m/m.y. The sedimentbasin in regionB is Determiningthe compensationmechanism of the Tuamotu boundedon the westby a scarpextending from Mataiva and to Plateauis importantin constrainingthe elasticstrength of the the eastby a nearlyparallel scarp extending northwest from lithosphere,which reflects the ageof theunderlying plate at the Tikehau. The total sediment accumulation for this basin is -1.6 time of accretion. The compensationmechanism also is km,suggesting an averagesedimentation rate of-32 rn/m.y.The importantin determiningthe volcanicthickness of the plateau sedimentationrates for thesebasins and for the centralportion of and thus constrainingthe volumeand flux of eruption,which theplateau are comparable with those at DSDPsite 318 which reflect the nature of the mantle source. The interface between the rangefrom 44 rn/m.y.,during the lower Pleiocene, to 6 rn/m.y. baseof the crust and mantle (Moho) was imagedwith reflection duringthe lower Miocene. dataonly outside of theplateau margins, restricting constraints on The sedimentbasin in regionC is boundedon the westby a the depthand shapeof the Moho to the sidesof the plateau. scarpalong the northside of a smallseamount (seamount A, Therefore, to constrainthe deep crustalstructure beneath the Figure2). Volcanicbasement in regionC maylie asdeep as 5.0 platform,we focuson the gravity and bathymetry which are both km below the present-dayseafloor, correspondingto the sensitive to variations in crustal thickness. predicteddepth of thedeflected, preexisting seafloor upon which theplateau formed (from lithospheric flexural models constrained Crustal Compensation by thegravity data as discussed later). Thus, the base of thisdeep valleymay have never been covered by Tuamotuvolcanism. The Here we assumethat the Tuamotu magmasloaded the upper averagesedimentation rate in regionC is as muchas -100 surfaceof the preexistingseafloor, thereby downwarping the rn/m.y.,several times that of therest of theplatform. Offsets in lithosphericplate. Thus, topography of theplateau is isostatically sedimentstratigraphy through layer 1 andin theupper section of supportedby thethickened crust, and the free-air gravity is the layer 2 indicatefaulting, most likely from slumpingof the combined attraction of the seafloor-water interface and the Moho. sedimentarylayer downthe slopeof seamountA. The more Given the observedtopography, we makepredictions as to the subduedtopography near the northeastern margin compared with shapeof theMoho for differentelastic-plate thicknesses (Te). We that of the southwesternmargin may indicate that gravity thenproduce theoretical gravity profiles that we compare with the slumpingwas more prevalentin the northeast.This might observedfree-air gravity data. Since the gravity signal is explainthe much higher sedimentation rate for basin C relativeto sensitiveto three-dimensionaldensity structure, we performthe the restof the platform. flexure and gravity calculationsusing the griddedbathymetry Thedeep reflections, marked R2 (Figures3b and3c) begin at a mapbut compare only those values coinciding with our ship track depthof-4.7 km near212øE and deepen to almost8.0 km near to the observedgravity. 212.5øE,where they reach-6 km below the plateau'ssurface. A standard method of solving the 2-D, fourth-order, These reflections are from a deep crustal impedancecontrast, differentialequation describing the deflectionof an elasticplate possiblythe upper surface of the gabbroiclayer underlying the loaded from above, is to solve the equation in the spectral basaltsas suggestedby Talandierand Okal [ 1987]. Talandier domain. The solution of the 2-D transformof downwardplate andOkal find a velocityof 6.83krn/s for thisdeep layer, which is deflectionW dependson the 2-D transformof topographyH, ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8105 accordingto o- œr3 (2) W--Apc 1q- (211;k)4•D H, (1) Apm (Apm)g whereE is Young'smodulus (8x10 •ø N/m2) andv is Poisson's ratio(0.25). UsingParker's [ 1973]spectral method, we calculate in which Apc and Apm are the crust-water and mantle-crust the combinedgravity fields due to the 2-D topographyand the density contrastsrespectively, k is the amplitudeof the 2-D deflectionof the Moho. We then producedtheoretical free-air wavenumber,and g is the accelerationof gravity. The flexural gravitygrids from which we extractedvalues coinciding with our rigidityD dependson elasticplate thickness Te, accordingto shiptrack. The densityof the waterand mantle are assumedto be

o5 80-1 •okm ......

40• •?•'.,",: ,,,,':•)•_/%.d&.•:. :',,.•",,,.,.. •. 0kan...... 0 " ' "2.'> •

-40

, 210 211 212 213 214 Longitude b) 2900 i i

2800

2700

2600

2500 0 5 10 15 20 25 Te (km) c) 80- Pratt isostasy 40-

-40 -

I I I I 210 211 212 213 214 Longitude Figure 6. (a) The depthprofile alongthe shiptrack lies abovethe observedfree-air gravity profile (thick, gray- shaded)and theoretical profiles assuming an averagecrustal density of 2650 kg/m3 and elastic plate thicknesses (Te) of 0, 5, and 10 km. The elasticplate model for T.•=5km yieldsthe smalleststandard deviation misfit of 5.9 mGal, while the modelsfor Te=Oand 10 km yield misfitsof 9.7 and 11.9 mGal respectively. (b) Standard deviationmisfits contoured at 2-mGalintervals for gravitymodels of crustaldensities, 2500-2900 kg/m 3 and elasticplate thicknesses 0-25 km. (c) The observedand theoretical gravity profile assuming Pratt isostasy with a depthof compensation,Zc, of 20 km. The theoreticalgravity includes the attractiondue to the seafloor-water densityinterface; a crust-mantle interface assuming the crust is a constant6 km in thicknessand 2560 kg/m 3 in density;and lateral density variations in the mantle,Ap, setaccording Pratt's equation: Ap=poH/z•., where Po is themantle-water density contrast (2300 kg/m 3) andH is topography. 8106 ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

1000 and 3300 kg/m3 respectivelyand the two parameterswe contributessignificantly to supportingthe topographyof the vary are Te andApc. Walvis Ridge in the SouthAtlantic. The theoreticalgravity Theoreticalprofiles for elasticplate thicknesses 0, 5, and10 profilein Figure6c is thatpredicted by assumingthe topography km, andcrustal density 2650 kg/m 3 arecompared with observed of theplateau is compensatedby lateraldensity variations in the free-airgravity profiles in Figure6a. The5-km-thick elastic plate mantle,20 km below a crustwith a uniformthickness of 6 km. modelyields the lowest standard deviation misfit (5.9 mGal) to The profileoverestimates the amplitudeof theobserved gravity the datafor this crustaldensity. Standarddeviation misfits for anomaly over the central portion of the plateau and elasticplate thicknesses ranging 0-25 km andfor average crustal underestimatesthe amplitudesto the sidesof theplatform which densities2500-2900 kg/m 3 arecontoured in Figure6b. For all come from the flexed Moho. The standard-deviation misfit for densities,misfits increase dramatically away from T•= 5 km. The thisprofile is 12 mGalbecoming larger with increasing assumed maximumof 44.0 mGal is obtainedfor T•=25 km and crustal depthsof compensation.Thus our gravity modeling shows that density2900 kg/m 3. Thelowest misfits (<6.0 mGal) are found the dominantsource of compensationis low-densitymaterial for densities2550-2750 kg/m3 andT•=5 km. Sinceall crustal within or near the base of the crust. densitiesexamined yield minimummisfits near T•=5 km, we concludethat the effective elasticplate thicknessbeneath the DeepCrustal Structure and VolcanicProduction northwesternportion of theTuamotu platform is 5+5km. This As noted above, Moho reflectionswere observedin the MCS thicknessis similar to elastic plate thicknessesbeneath the datato the sidesof the plateaubeneath the normaloceanic crust southernend of the plateauof 2-6 km, obtainedfrom the geoid (Figure7). Thesearrivals, while low in amplitude,are prevalent analysisof Okaland Cazenave[1985]. Ourlowest standard throughoutthese sections, consistently -2 s belowthe volcanic deviationmisfits (of which1-2 mGalare the intrinsic error of the basementwhich is blanketedby 0.2-0.5 s of pelagicsediments. gravimeter)are higher than the 2-5 mGalmisfits obtained by Usinga velocityof 6 km/sfor the volcaniccrust as constrained Filmeret al. [ 1993]over the Marquesas and Society Islands using from the semblancedata and refraction data from sonobuoy3 thesame technique, most likely due to theunmodeled 1 to 2-km- launched near 214øE (see Figure 2), we constrainthe local thicksediment layer present over Tuamotu volcanics but absent thickness of the oceanic crust to be 6+1 km. overthe Marquesas and Society volcanics. By incorporatingthese constraints for thecrust to thesides of the plateauwith the deep crustalstructure from the gravity Pratt Compensation modelingand the sedimentarystructure over the platform,we While crustalcompensation is clearlyimportant, another summarizeour crustalmodel in Figure8. The shapeof the Moho sourcethat may contribute to thecompensation of the Tuamotu as derivedfrom our best fitting flexural model (T•=5 km and Plateauis lateraldensity variation in the lithosphere(Pratt-type crustaldensity 2650 kg/m3, solidcurve) is consistentwith the isostasy)that may have resulted from varying degrees of partial seismicobservations to the sidesof the Tuamotu platform. We meltingof themantle during the plateau's accretion. Angevine showthe top of the preexistingseafloor (dashed) assuming that andTurcotte [ 1980]suggested that this compensation mechanism all Tuamotuvolcanism was emplacedon top of the 6-km-thick

5 Southwest' '••ll•

••...•:'4..- ;,•:.:• •. ,•,• •:,• •..,•,• •.,.,• .,,.•':... •,. '-.,: :.--., x•,...,..•. ,:,.,.• ...... •..x•.• .• ,•,.

5 Northeast

.:.• ....

Figure7. Timesections of theabyssal seafloor to the southwest and northeast of the Tuamotu Plateau. Moho arrivalsare observed consistently -2 s belowthe sediment-volcanic basement interface (marked b). ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8107

• , 12 ", .-'

14 Moho Reflectors

18 l 210.5 211 212 213 214 215 Longitude Figure 8. Depth profile alongthe shiptrack showingsedimenta• layer 1, the estimatedlocation of the volcanic basement(shaded), Moho reflectionsin the surroundingseafloor, and the theoreticalMoho profile of our best fitting elasticplate model (solid). •e dashedline m•ks the preexistingseafloor assuming the Tuamomplatform loadedthe uppersurface of a unifo•, 6-km-thickoceanic crest.

oceanic crust. Note that the base of region C, lies near the and subsurfacevolumes assuming topography is compensatedby preexistingseafloor, indicating as mentionedpreviously, that Airy isostasy[Schubert and Sandwell, 1989]. The volume there may have been very little excessvolcanism in this basin. estimate for Hawaii, considers only the volume of surface The R2 reflections lie 3-9 km above the preexisting seafloor, topographyof the chainfrom the EmperorSeamounts to the midwaybetween the preexistingseafloor and sediment-basement island of Hawaii (it does not include the compensatingcrustal interface. We concludethat the maximumcrustal thickness along root, but this may be offsetby the fact that it doesinclude the our profile is -18 km, the maximumtotal volcanicthickness is topographicswell not associatewith crustalthickening). The 16-17 km, and the maximum excess volcanic thickness (i.e., in volume estimate for the MarquesasIslands from Wolfe et al. addition to a 6-km-thick normal oceaniccrust) is 9-10 km. [1994] includes the volumes from the topography and From these constraints we can estimate the total volcanic compensating roots as well as the volume of the volcanic production of the Tuamotu hotspot. Using the gridded sedimentsfilling the surroundingflexural moat as constrained bathymetry map of the region, we sum the volume from the seismically. The excess volcanic volume of the Tuamotu topography (ie., shallower than the average depth of the platform is significantlylarger than the volume of the Marquesas surroundingseafloor 4.2 km), as well as the volume of the Islands; and although the Tuamotu Plateau is morphologically compensatingcrustal root, assumingthe entire platform is distinctfrom Hawaii, the excessTuamotu volume is comparable compensatedby an elasticplate thicknessof 5 km. From this to that of Hawaii. The Tuamotu Plateauis significantlysmaller volume we subtractthe sedimentvolume, assuminga sediment than the otherplateaus despite its morphologicalsimilarity. thicknessof 0.2-0.3 km beneathseafloor deeper than 3.5 km and Possibly,more telling as to the natureof the mantlesource of 1-2 km beneath seafloor shallower than 3.5 km. Our estimate for these volcanic features is the rate at which these volumes were thetotal volcanic volume of theplateau is 7.3-7.8x106km 3 and extruded(Figure 9b). We derive the volcanicfluxes for Manihiki our estimatefor the excessvolcanic volume (ie., subtractingthe and Ontong Java Plateaus (1.7-4.4 and 7.9-16.9 km3/yr volumeof a 6-km-thicklayer) is 2.0-2.6x106km 3. If we assume respectively)by dividing the volumesby the 3-m.y. durationof a fixed hotspot,the Pacific plate took 20 m.y. to migrate along volcanismas estimatedfrom drill core findings [Tarduno et al., the length of the platform (from absolutePacific plate rotation 1991; Coffin and Eldholm, 1993], and flux ratesfor BrokenRidge rates of Duncan and Clague [1985]). This time span for the and Kerguelen(1.2-2.8 and 5.5-9.9 km3/yr respectively)by plateau'sformation yields averagevolcanic fluxes, including and dividing by the volcanicduration of 4.5 m.y. alsofrom drill core excluding6 km of thecrust, of-0.4 and0.1 km3/yrrespectively. data [Coffin and Eldholm, 1993]. Volcanism along the Walvis The volume and fluxes which include 6 km of the crust may be Ridge, between the continental margin of Africa and -36øS, the most appropriateestimates if the plateau formed on-axis, ensued for -70 m.y. according radiometric dating results of while those which exclude the normal oceanic crust may be the O'Connor and Duncan [1990] which yields an averageeruption mostappropriate, if the plateauformed off-axis. rate of 0.1-0.2 kmS/yr. Eruptionfluxes for Hawaii and the To characterize the mantle source which gave rise to the MarquesasIslands (0.5 and0.4 km3/yrrespectively) as estimated TuamotuPlateau, we comparevolcanic volumes for the platform by Filmer et al. [1994], are the averagefluxes over the past 6 with those of two island chains and five oceanic plateaus in m.y.; they includetopographic and compensatingvolumes as well Figure9a. The volumesof the five plateausinclude the surface as the volumes of the sediment-filled moats. The flux estimates 8108 ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

a) b) s0 15 • 40

• 30 • lO

> 20 ¸

¸ o•

Figure 9. Bar graphsof the estimated(a) eruptionvolumes and (b) eruptionrates of the Hawaiianand Marquesas Islandchains, the TuamotuPlateau, and five oceanicplateaus. White hachuredbars denote estimations using the total volcanicvolumes; black bars denote values using only the excessvolcanic volumes (i.e., totalvolume minus the volume of the normal oceaniccrest: 6 km in thicknessas constrainedin this studyfor the TuamotuPlateau, and 6.5 km in thicknessas used by Shubertand Sandwell[ 1989] for the otheredifices). The uncertaintyrange for the Tuamotu total and excessvolumes are the unpattemedsections of the bar.

for the Tuamotu Plateau are most comparablewith thoseof [Schlangeret al., 1984]. Thus, accretionof the northernportion Hawaii and Marquesasand, with the exceptionof the Walvis of the Tuamotu Plateau may have begun as early as 55 Ma and Ridge,substantially less than the otheroceanic plateaus. Again, persistedto at least41.8 Ma. the mantlesource which generated Tuamotu magma is similarin When consideringthe results of a recent magnetic study by both size and longevity to the hotspotswhich have formed Antoine et al. [1993], the above ages suggestthat the northern Hawaii andthe Marquesas.The WalvisRidge source, although portionof the Tuamotuplatform formed significantly off the axis liberatinga greatermagmatic volume, appears to be a longlived of the Pacific-Farallonspreading center. As illustratedin Figure hotspotsimilar to the Tuamotu hotspot. The other oceanic 10, the sectionof the plateaunear our surveyline is undefiain by plateausappear to haveformed by a very differentphenomena oceaniccrust with ages -64-71 m.y.; therefore,an age of 50-55 which involved dramatic flood-basalt volcanism. m.y. for the northernplateau suggeststhat it was eraplacedon lithosphereof age -10-20 m.y. The spreadingcenter at this time Implications for the Origin of the Tuamotu was approximately600 km to the eastof the plateau(Figure 10). Plateau A lithosphericplate with age 10-20 m.y. would have an effective elastic thicknessof 10-15 kin, assumingT e correspondsto the Proximity to the Pacific-Farallon Spreading Center depthof the 450øC lithosphericisotherm [Watts, 1978; Wattset If the mantle sourceof Tuamotu volcanismis a hotspotnmch al., 1980]. This range of thicknessis slightly greater than our like the Hawaiian and Marquesashotspots, we mustnow explain upper bound of 10 km obtained from our gravity modeling why the continuousvolcanic ridges and plateaulike form of the possiblybecause the regional lithospheremay be anomalously TuamotuPlateau differ from the discretepinnacles that compose warm [McNutt and Fischer, 1987], thusyielding extremelysmall the two island chains. To better understandthe magmaticand valuesof Te [Menardand McNutt, 1982;Calmant and Cazenave, tectonicprocesses which constructedthe northwesternend of the 1987]. Errors in the geomagnetictimescale of Berggren et al. Tuamotu Plateau, we nmst first considerthe tectonicsetting in [ 1985] are likely to be lessthan 2-3 m.y. (seeFigure 30 of Cande whichthe plateauformed. The proximityof the TuamotuPlateau and Kent [1992]); therefore, unless the age constraints for to the Pacific-Farallon paleo-spreadingcenter (at which the Tuamotu volcanism are significantly in error, it is unlikely that underlyingoceanic lithosphere was created)can be inferredfrom the northwestern end of the Tuamotu Plateau formed at the available age constraints for the plateau compared with Pacific-Farallonpaleo-ridge. geomagneticages of the surroundingseafloor. The n•inin•umage The inferred subsidenceof the plateauis also consistentwith of the northwesternend of the platformis constrainedby the age the 10-20-m.y. age difference between the plateau and the of the oldest sediments sampled on the plateau (Figure 10): underlyinglithosphere. Isostaticremoval of the 1-2 kin-thick lower middle Eocene,from a dredgehaul alongthe southwestern sedimentlayer (using an average sediment density of 2100kg/m 3 margin near 14.5øS [Burckle and Saito, 1966]; lower Eocene, [Schlangeret al., 1976]) would put the basementdepth over the from DSDP site 318 [Martini, 1976]; and middle Eocene, from central portion of the plateau at 1.6-2.1 kin. This depth range dredgehauls along the southwesternmargin near 16.5øS[Le may reflect the subsidenceof the plateauif it was oncenear sea Suaveet al., 1989]. Volcanicsamples from the plateauare from level, as suggestedby the reefal sedimentsobtained at DSDP site two dredgehauls along the southwesternmargin very nearour 318. Sucha rangeof subsidenceis expectedif the plateauloaded surveyline with 4øAr?Ar datesof 41.8_+0.9Ma and47.4_+0.9 Ma a coolinglithospheric half-space of age5-14 m.y., consistentwith ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8109

% \ \ \ 21 20 18 16 g•.•\ [ (39.2) _15 ø •:.::::. 4:::::::::

......

a- Lower Middle Eocene (44-51 Ma) [Burckle and Saito, 1966] b- 40.9-48.3 Ma [Schlangeret al., 1984] c- Lower Eocene (52-55 Ma) [Martini, 1976] d- Middle Eocene (40-51 Ma) [Le Suave et al., 1989]

Figure 10. Map of magneticlineation anomalies[Antoine et al., 1993] surroundingthe Tuamotu Plateau (shadedfor depths<3500 m) marked with anomalynumbers and correspondingages (Ma) [Berggrenet al., 1985]. The configurationof the paleo-spreadingcenter near the time of formationof the plateauis shownby the thickestisochron labeled with its age, 50.3 m.y. Locationsof sedimentand crustalsamples are labeledaccording to their agesand sourcesas listed. Our shiptrack is markedby the thickline crossingthe northwesternend of the platform.

our age estimate of 10-20 m.y. The predicted subsidence, marginformed either along a paleo-transformfault separatingthe assuminga plate age of 10-20 m.y. at the time of loading,is 1.8- Farallon and Antarctic plates, or along the paleo-spreadingaxis 1.4 km, much lessthan the -3 km of subsidencepredicted if the separatingthe Antarcticand Pacific plates. A recentmagnetic plateauformed on axis. and gravity surveyto the east of the plateauby Munch et al. [ 1992] supportsWinterer et al.'s laterhypothesis. Morphological Similarities: The Manihiki Plateau The Walvis Ridge is also similar to the TuamotuPlateau in and the Walvis Ridge manyaspects. It is borderedto the northby a volcanicridge and is composedof numerousvolcanic ridges. The profile in Figure Although the ages for the platform and underlying seafloor llb, obtained from the seismic profile of Goslin and Sibuet indicatean off-axis origin for the Tuamotuplatform, the volcanic [1975], illustrates the northern margin of the Walvis Ridge, morphology is similar to that of many oceanic plateaus that which stands >2 km above the seafloor and dams the plateau formed at oceanic spreadingcenters. When consideredwith the sediments. The Walvis Ridge is also similar to the Tuamotu surroundingbathymetry, the volcanicscarps and moundsimaged platformin the asymmetryof their correspondingprofiles: the by the reflection seismics appear to be linear features with abrupt northernmargin and gradual tapering of the southern roughlynorthwest to southeastlineaments (Figure 2). Despitethe margin of the Walvis Ridge resemblesthe southwesternand numerousseamounts and atolls scatteredthroughout the Tuamotu northeasternmargins of the Tuamotu Plateau, respectively. platform, the dominant crustal fabric appearsto be continuous, Goslin and Sibuet[ 1975] suggestthat the Walvis Ridge formed elongated volcanic features quite unlike what is found within during or near the initial breakup of the African and South midplateisland chains. American continents,and that the northernWalvis margin marks Two oceanic plateaus with ridges and scarpssimilar to the the boundaryof a paleo-transformfault. Tuamotu platform are the Manihiki Plateau, located -1000 km northwest of the Tuamotu Plateau (Figure l a), and the Walvis Model for the Formation of the Tuamotu Platform Ridge which extendssouthwest from the continentalmargin of Africa into the southernAtlantic. The basementrim along the Thus, when consideringthe Manihiki Plateau and Walvis eastern margin of the Manihiki Plateau [Winterer et al., 1974] Ridge examples,it is likely that the formation of the linear bears close resemblance, in both shape and scale, to the volcanic features of the Tuamotu Plateau was influenced by southwesternmarginal ridge of the Tuamotu platform (Figure tectonic structures in the underlying crust and lithosphere. 1l a). The Manihiki marginal ridge stands -3.4 km above the Structuresrelated to rifting at an oceanspreading center as in the seafloorof the Penrhyn Basin and, like the Tuamotu margin, case of the Manihiki Plateau can be excluded, since the Tuamotu forms a barrier bounding the plateau sediments. Ridges and platform was formed significantly off-axis and the plateau faults observedin seismic profiles along the easternmargin of lineamentsare oblique to seafloorisochrons. Likewise, it is Manihiki Plateauled Winterer et al. [ 1974] to hypothesisthat this unlikely that the Tuamotu platform formed along a transform 811o ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

a) 1oøs

2 3 4

6 7 8

Figure 11. (a) Stratigraphicprofile (from Wintereret al. [1974]) of the Manihiki Plateaualong the shiptrack markedby arrowsin the adjacentmap (the directionof the arrowsdenote left-m-right in the profile). The plateau sedimentarylayers are shownoverlying volcanic basement (gray-shaded). Note the volcanicmounds composing the east (marked by the arrow) and west margins resemble volcanic mounds of the Tuamotu Plateau. The seafloor of the Tuamotu platform (shaded profile) is shown at the same horizontal and vertical scale for comparison.(b) Stratigraphicprofile of the Walvis Ridge alongthe profile markedby the arrow in the adjacent map (1-km contourinterval with depthsshallower than 4 km shaded),drawn at the samehorizontal and vertical scale as the Manihiki profile. Again the volcanicbasement is gray-shaded,the overlying sedimentlayers are marked, and the seafloorof the Tuamotu platform is shownat the samescale. The arrows above the profiles mark the northernmarginal ridge. zone, as proposedfor the Walvis Ridge, sincethe lineamentsof lithospheresufficient to channelhotspot melts, thereby generating the Tuamotuvolcanic ridges, and the overallplatform trend, are the linear volcanicridges of the plateauand hidingthe changein oblique to the nearby Marquesasand Austral fracturezones. Pacific plate motion at 42 Ma. Outside of this lithospheric Still, the plateaulineament does not directlyreflect the motionof transferzone, regionssuch as the basin in region C may have thePacific plate over the Tuamotu hotspot, since (assuming this experienced less volcanism owing to the more competent hotspotwas stationarywith respectto the Hawaiianhotspot) the unscarredlithosphere. changein Pacificplate motion at -42 Ma [Duncanand Clague, Lithosphericandcrustal control on volcanic morphology is 1985; Lonsdale, 1988] is not recordedin the Tuamotulineament. discussedby Vogt and Johnson[1975], who show that plate A plausibletectonic mechanism that would leave a lithospheric boundariessuch as transformfaults can channelhotspot-derived imprint alongthe Tuamotulineament is a southwardpropagating magmato form ridge like edifices. Examplesof suchvolcanic segmentof the Pacific-Farallonspreading center. Analogousto ridges are along the Clipperton, Galapagos,and Marquesas the Galapagospropagating rift tip near95.5øW [Hey et al. 1986], fracture zones. These ridges are thoughtto mark the relative the possibleTuamotu propagatormay have generatedan inner motion of the Marquesas hotspot as it passed beneath the pseudofault (the boundary between crust generated at the lithosphericdiscontinuities at thesefracture zones [McNutt et al., northernridge segment and that at the southern)and a failedrift 1989]. Accordingto Epp [!984], lineamentsin the Hawaiianand zone (the lineation of the failed southernspreading segment, Musician Island chainsdeviate from hotspottraces in responseto Figure 12). Thesetwo boundarieswould enclosea block of lithosphericzones of weaknesssuch as fracturezones. Direct lithospheretransferred from the Farallonplate to thePacific plate evidencefor channellingof hotspotmaterial toward lithospheric calledthe "lithospherictransfer zone" by Hey et al. [ 1986]. The discontinuitiesis the isotopic and trace element enrichments inner pseudofault,the failed rift zone, and the lithospheric found along midoceanridge segmentslocated near hotspots transfer zone may have been weaknessesin the crust and [Schilling, 1985; Schilling et al., 1985]. These enrichments ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8111

50 Ma

_15 ø

\ ,

_20 ø P \ Farallon Plate

210 ø 215 ø 220 ø 225 ø Longitude

Figure 12. This cartoonillustrates the geometryof the Pacific-Farallonspreading center (thickest lines) at time 50 Ma. Isochrons(extrapolated to inferredplate boundaries)with agesgreater than 50 Ma are solid, while those with ages less than 50 Ma are dashed. Our preferred model that the northern ridge segment propagated southwardat the expenseof the southernsegment predicts a zone of lithospheretransferred from the Farallon plate to the Pacific plate (hachured)bounded to the northby the inner pseudofaultand to the southby the failed rift zone. These discontinuitiesin the lithospheremay have focused volcanism along the Tuamotu Plateau (outlinedwith the 3.5-km depthcontour) and shapeits plateaulike morphology.

indicate that hotspot material may deviate laterally from the Ducie seamount chain just south of the Tuamotu Plateau; hotspotcenter by as much as 800 km as a consequence,at leastin however,in our model, this hotspotwould havebeen too far west part, of the thinnedlithosphere at midoceanridges. to have contributed to the Tuamotu Plateau. Our model is The most direct evidence for a lithospheric suture zone consistentwith the model of Okal and Cazenave [ 1985] in that delineating the path of a propagating rift tip of the Pacific- the Sala y Gomez hotspot would have been on axis when it Farallon spreading center is seen in the regional magnetic generatedthe southeasternend of the plateau. We based our isochronpattern (Figure 10). The magnetizationpattern reveals model on the rotationpoles of Duncan and Clague [ 1985], which an age offset betweencrust on the southwestand northwestof the are constrainedby the age progressionalong the Hawaiian chain. platform; and beginningwith anomaly20, the isochronsnorth of We alsoassume that all of the hotspotsare stationarywith respect the plateaulengthen to the southwith decreasingage, indicating to the Hawaiian hotspots. that the northernsegment was propagatingsouthward when these A secondpossibility is basedon rotation poles of Lonsdale isochronswere formed. This is the same propagatingrift that [1988], which also fit the data from Hawaii, but are derived from Okal and Cazenave [1985] suggest in their model for the volcanicages along the Louisville Ridge in the SouthwestPacific formation of the southeasternend of the Tuamotu platform. Basin. In this model (Figure 13b) the Salay Gomezhotspot may While the magnetic data show clear evidence for an offset have been well east of the Tuamotu platform at-50 Ma, while betweenthe northernand southernridge segments,whether the hotspot"X" would be in closerproximity to the northwesternend northernsegment was propagating before the time of anomaly20 of the plateau. In this model, hotspot X generated the is less certain. The propagatingrift model, however, is most northwesternend of the plateau, while and the Easter hotspot attractivein that it may explain the morphologyand geographic generated the central and eastern portions. This model is trend of the Tuamotu Plateau. Thus, while the mantle source that consistentwith the conclusionsof Woodsand Okal [ 1994], based gave rise to the TuamotuPlateau is most akin to the hotspotsof on crustal structure arguments, that Sala y Gomez is not island chains, the morphology of the plateau appearsto be associated with the formation of the Tuamotu Plateau or its structurallycontrolled by the lithosphere,similar to other oceanic Nazca-plateconjugate, the Nazca Ridge. plateaus. If we assumehotspots are stationarywith respectto eachother, As for the hotspotswhich formedthe Tuamotuplatform, one the models above suggestthat the Tuamotu Plateau was most possibility,is that the platformwas formedby the Sala y Gomez likely formed by two hotspots,one of which was the Easter and the Easterhotspots. As illustratedin Figure 13a the hotspot, hotspot.These hotspots liberated magma with a flux comparable now beneathEaster Island, may have been beneaththe plateau to those of the Hawaiian and Marquesas hotspots, but near our survey line at 50 Ma, while the hotspotnow beneath substantiallyless than thosethat gave rise to the flood basaltsof Sala y Gomez, may have beenbeneath the southeasternend of the other oceanicplateaus. Tuamotu Plateau at or near the Pacific-Farallonspreading center. The relative path of the Sala y Gomez hotspotmay explain the Conclusions eastwardjog of the Tuamotuplatform near 217øE. Additionally, as proposedby Okal and Cazenave [1985] another hotspot We have developed a scenario for the formation of the (marked X, Figure 13a) may have formed the Acteon-Oeno- northwestern end of the Tuamotu Plateau. The volcanic 8112 ITO ET AL.' CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

a) _10 ø

_30 ø

b) _10o

-200

_30 ø 210 ø 220 ø 230 ø 240 ø 250 ø 260 ø Longitude Figure13. Theconfiguration of the Pacific-Farallon spreading center at 50Ma is shownas the thickest solid lines.The large dots mark the locations ofthe Sala y Gomezhotspot (S), the Easter hotspot (E), and the unknown hotspot(X, suggestedbyOkal and Cazenave [1985]) at -50 Maderived from Pacific plate rotation poles of (a) Duncanand Clague [1985] and (b) Lonsdale [1988]. The migration path of thesehotspots with respect to the Pacificplate is shownas thick shaded lines with crosses marking the locations at10-m.y. intervals.

constructstake the form of linear, continuousridges and scarps Plateau sets this hotspot feature apart from typical oceanic which bound a carbonateand volcaniclasticsediment layer with plateausand island chains. thickness 1-2 km over the central portion of the plateau. Radiometric and paleontological ages of the plateau and Acknowledgements.The success of thecruise EW9103 was due largely geomagneticages of the seafloorindicate that the northwestern to efforts of our co-chief scientist John Mutter, as well as the onboard endof the plateauformed off-axis by -600 km, on lithosphereof scientific and technicalstaff, and the crew of the R/V Maurice Ewing. age 10-20 m.y. An off-axisorigin is consistentwith our upper We thankNathan Bangs for the useof his semblanceanalysis software, PeterBuhl for histime and advice put forth in thepreliminary stacking of boundestimates for the elasticplate thickness of 5-10 km derived the MCS data,and Marc Munschyfor supplyingthe coordinatesof his from the gravity and bathymetrydata, and with the subsidence seafloormagnetic lineation picks. Finally, we thank Carol Raymond, Uri implied by isostaticremoval of the plateausediments. We S. Ten Brink, andDavid Sandwellfor their constructivereviews of this attributethe formation of the northwesternend of the plateauto manuscript.This project was funded by NSF,grant OCE-8817764. the Easter,and one otherhotspot, which eruptedmagmas along the suturezone of a southwardpropagating rift segmentof the References Pacific-Farallonspreading center approximately 50 Ma. This Angevine,C. L., and D. L. Turcotte,On the compensationmechanism of suture zone channelled hotspot melts to shape the volcanic the Walvis Ridge, Geophys.Res. Lett., 7, 477-479, 1980. morphologyquite differentlyfrom that of islandchains formed Antoine,C., M. Munschy,and A. Gachon,Structure and evolutionof the by mantlehotspots of comparablesize and strength.While the oceaniccrust in the southof Tuamotu Islandsregion (South Pacific), local lithosphericstructure may have controlledthe plateaulike paper presented at the International Workshop on lntraplate morphologymuch like otheroceanic plateaus which formed near Volcanism: The PolynesianPlume Province, Papeete,Tahiti, Aug., plateboundaries, the mantlesource may havedetermined the 1993. magmaticflux which is comparableto the Hawaiian and Berggren, W. A., D. V. Kent, J. J. Flynn, and J. A. Van Couvering, MarquesasIsland chains. This uniqueorigin for the Tuamotu Cenozoicgeochronology, Geol. Soc.Am. Bull., 96, 1407-1418, 1985. ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU 8113

Boyce,R.E., Soundvelocity-density parameters of sedimentand rock Central Pacific Ocean (DSDP Leg 33), Initial Rep. Deep Sea Drill. from DSDP •drill sites 315-318 on the Line Island chain, Manihiki Proj., 33, 383-397, 1976. Plateau,and Tuamotu Ridge in the Pacific Ocean,Initial Rep. Deep McNutt, M. K., and K. Fisher, The South Pacific superswell, in SeaDrill. Proj. 33, 695-728, 1976. Seamounts,Islands, and Atolls, Geophys.Monogr. Ser., vol. 43, edited Burckle, L. H., and T. Saito, An Eocenedredge haul from the Tuamotu by B.H. Keating,P. Fryer, R. Batiza, and G. W. Boehlert,pp. 25-34, Ridge,Deep Sea Res. Short Commun., 13, 1207-1208,1966. AGU, Washington,D.C., 1987. Calmant,S., and A. Cazenave,Anomalous elastic thickness of the oceanic McNutt, M. K., K. Fischer,S. Kruse,and J. Natland,The origin of the lithospherein thesouth-central Pacific, Nature, 328, 236-238,1987. Marquesasfracture zone ridge andits implicationsfor the natureof hot Cande,S. C., and D. V. Kent, A new geomagneticpolarity time scalefor spots,Earth Planet.Sci. Lett., 91,381-393, 1989. the Late Cretaceousand Cenozoic, J. Geophys.Res., 97, 13,917- Menard, H. W., and M. K. McNutt, Evidence for and consequencesof 13,951, 1992. thermalrejuvenation, J. Geophys.Res., 87, 8570-8580,1982. Clague,D. A., andR. D. Jarrard,Tertiary Pacific plate motions deduced Morgan,W. J., Plate motionsand deepmantle convection, Mern. Geol. from the Hawaiian-EmperorChain, Geol. Soc. Am. Bull., 84, 1135- Soc. Am. 132, 7-22, 1972. 1154, 1973. Morgan,W. J., Hotspottracks and the openingof the Atlanticand Indian Coffin, M. F., and O. Eldholm, Scratchingthe surface: Estimating Oceans, in The Sea, vol. 7, The Oceanic Lithosphere, edited by C. dimensionsof largeigneous provinces, Geology, 21, 515-518,1993. Emiliani, pp. 443-487, JohnWiley, New York, 1981. Detrick, R. S., and A. B. Watts, An analysisof isostasyin the world's Munch, C., S. Kruse, and M. McNutt, Hydrosweep,gravity and magnetic oceans,3., Aseismicridges, J. Geophys.Res., 84, 3637-3653,1979. data across an abandoned rift southeast of the Manihiki Plateau Detrick, R. S., J. G. Sclater,and J. Theide, The subsidenceof aseismic (abstract),Eos Trans.AGU., 73 (43), Fall Meeting suppl.,585, 1992. ridges,Earth Planet. Sci. Lett., 34, 185-196,1977. Munschy,M., and R. Schlich,Structure and evolutionof the Kerguelen- Duncan,R. A., andD. A. Clague,Pacific plate motion recorded by linear Heard Plateau (indian Ocean) deduced from seismic stratigraphy volcanicchains, in The OceanBasins and Margins, editedby A. E. M. studies,Mar. Geol., 76, 131-152, 1987. Nairn, F. G. Stehli, and S. Uyeda, pp. 89-121, Plenum,New York, O'Connor, J. M., and R. A. Duncan, Evolution of the Walvis Ridge-Rio 1985. Grande Rise hot spot system: Implicationsfor African and South Duncan,R. A., and I. McDougall,Linear volcanism in FrenchPolynesia, Americanplate motionsover plumes,J. Geophys.Res., 95, 17,475- J. Volcanol. Geotherm. Res., 1, 197-227, 1976. 17,502, 1990. Epp,D. Possibleperturbations to hotspottraces and implications for the Officer, C. B. and P. C. Wuenschel,Reduction of deepsea refraction data, originand structure of the Line Islands,J. Geophys.Res., 89, 11,273- Tech.Rep. 1, LamontGeol. Obs.,Palisades, N.Y., 1951. 11,286, 1984. Okal, E. A., and R. Batiza, Hotspots: The first 25 years,in Seamounts, Filmer, P. E., M. K. McNutt, and C. J. Wolfe, Elastic thicknessof the Islands, and Atolls, Geophys.Monogr. Ser., vol. 43, edited by B. lithospherein theMarquesas and Society Islands, J. Geophys.Res., 98, Keating and R. Batiza, pp. 1-11, AGU, Washington,D.C., 1987. 19,565-19,577, 1993. Okal, E. A., and A. Cazenave,A model for the plate tectonicevolution of Filmer, P. E., M. K. McNutt, and H. F. Webb, Volcanism and the east-centralPacific based on Seasatinvestigations, Earth Planet. archipelagicaprons in the Marquesasand HawaiianIslands, Mar. Sci. Lett., 72, 99-116, 1985. Geophys.Res., 16, 385-406, 1994. Parker,R. L., The rapid calculationof potentialanomalies, Geophys. J.R. Goslin,J., andJ. C. Sibuet,Geophysical study of the easternmostWalvis Astron. Soc., 31,447-485, 1973. Ridge,South Atlantic: Deepstructure, Geol. Soc. Am. Bull., 86, 1713- Richards, M. A., R. A. Duncan, and V. E. Courtillot, Flood basalts and 1724, 1975. hotspottracks: plume headsand tails, Science,246, 103-107, 1989. Henderson,L. J., Motion of the Pacificplate relative to the hot spotssince Sandwell,D. T., and K. R. MacKenzie,Geoid heightversus topography Jurassicand modelsof the oceanicplateaus of the Farallonplate, Ph.D. for oceanic plateaus and swells, J. Geophys. Res., 94, 7403-7418, thesis,Northwest. Univ., Evanston,Ill., 1985. 1989. Hey, T., M. C. Kleinrock,S. P. Miller, T. M. Atwater,and R. C. Searle, Schilling,J.-G., Upper mantleheterogeneities and dynamics,Nature, 314, Sea Beam/Deep-Towinvestigation of an activeoceanic propagating 62-67, 1985. rift system,Galapagos 95.5øW, J. Geophys.Res., 91, 3369-3393, Schilling, J.-G., G. Thompson,R. Kingsley, and S. Humphris,Hotspot- 1986. migrating ridge interaction in the South Atlantic, Nature, 313, 187- Jackson,E. D., and S. O. Schlanger,Regional syntheses,Line Islands 191, 1985. Chain, Tuamotu Island Chain, and Manihiki Plateau, Central Pacific Schlanger,S. O., andthe ShipboardScientific Party, Site 318, Initial Rep. Ocean,Initial Rep. Deep Sea Drill. Proj., 33, 915-927, 1976. Deep Sea Drill. Proj. 33, 301-338, 1976. Jarrard, R. D. and D. A. Clague, Implications of Pacific island and Schlanger,S. O., M. O. Garcia, B. H. Keating, J. J. Naughton, W. W. seamountages for the origin of volcanicchains, Rev. Geophys., 15, Sager,J. A. Haggerty,J. A. Philpotts,and R. A. Duncan,Geology and 57-76, 1977. geochronologyof the Line Islands, J. Geophys.Res., 89, 11,261- Kogan, M. G., Gravity anomaliesand origin of the Walvis Ridge, J. 11,272, 1984. Geophys.Res., 84, 6019-6025, 1979. Schubert, G., and D. T. Sandwell, Crustal volumes of continents and of Larner, K., R. Chambers,M. Yang, W. Lynn, and W. Wai, Coherent oceanic and continentalsubmarine plateaus, Earth Planet. Sci. Lett., noisein marine seismicdata, Geophysics,48, 854-886, 1983. 92, 234-246, 1989. Larson, R. L., Latest pulse of Earth: Evidence for a mid-Cretaceous Talandier, J., and E. A. Okal, Crustal structure in the Society and superplume,Geology, 19, 547-550, 1991. Tuamotu Islands, French Polynesia, Geophys.J.R. Astron. Soc., 88, Le Suave, R., C. Pichocki, G. Pautot, M. Hoffert, Y. Morel, M. Voisset, 499-528, 1987. S. Monti, J. Amosse, and A. Kosakevitch, Geological and Tarduno, J. A., W. V. Sliter, L. Kroenke, M. Leckie, M. Mayer, J. J. mineralogical study of Co-rich ferromanganese crusts from a Mahoney, R. Musgrave, M. Storey, and E. L. Winterer, Rapid submergedatoll in the TuamotuArchipelago (French Polynesia), Mar. formationof OntongJava Plateau by Aptian mantleplume volcanism, Geol., 87, 227-247, 1989. Science, 254, 399-403, 1991. Lonsdale,P., Geographyand history of the Louisvillehotspot chain in the Vogt, P. R., Volcano spacing,fractures and thicknessof the lithosphere, SouthwestPacific, J. Geophys.Res., 93, 3078-3104, 1988. Earth Planet. Sci. Lett., 21,235-252, 1974. Luetgert,J. H., MacRay-interactivetwo-dimensional seismic raytracing Vogt, P. R., and G. L. Johnson,Transform faults and longitudinal flow for the Macintosh, U.S. Geol. Surv. Open File Rep., 92-356, 43 pp., below the midoceanridge, J. Geophys.Res., 80, 1399-1428, 1975. 1992. Watts, A. B., An analysisof isostasyin the world's oceans,1., Hawaiian- Martini, E., 9., Cretaceousto recentcalcareous nannoplankton from the Emperorseamount chain, J. Geophys.Res., 83, 5985-6005, 1978. 8114 ITO ET AL.: CRUSTAL STRUCTURE AND ORIGIN OF THE TUAMOTU PLATEAU

Watts, A. B., J. H. Bodine, and N.M. Ribe, Observationsof flexure and archipelagicapron: Seismicstratigraphy and implicationsfor volcanic geologicalevolution of the Pacific Oceanbasin, Nature, 283, 532-537, growth, masswasting, and crustalunderplating, J. Geophys.Res., 99, 1980. 13,591-13,608, 1994. Winterer,E. L., Bathymetryand regionaltectonic setting of the Line Woods, M. T., and E. A. Okal, The structureof the Nazca ridge and Sala Islandschain, Initial Rep.Deep Sea Drill. Proj., 33, 731-748,1976. y Gomez seamountchain from the dispersionof Rayleigh waves, Winterer, E. L., P. F. Lonsdale,J. L. Matthews,and B. R. Rosendahl, Geophys.J. Int., 117, 205-222, 1994. Structureand acoustic stratigraphy of theManihiki Plateau, Deep Sea Res., 21,793-814, 1974. R. L. Gibsonand M. McNutt, Departmentof Earth, Atmospheric,and Winterer,E. L., J. H. Natland,R. J. Van Waasbergen,R. A. Duncan,M. PlanetarySciences, Massachusetts Institute of Technology,Cambridge, K. McNutt, C. J. Wolfe, I. PremoliSilva, W. W. Sager,and W. V. MA 02139. Sliter, Cretaceousguyots in the NorthwestPacific: An overviewof G. Ito, Department of Geology and Geophysics, Woods Hole their geologyand geophysics,in The MesozoicPact.'l•c: Geology, Oceanographic Institution, Woods Hole, MA 02543 (e-mail: Tectonics,and Volcanism,Geophys. Monogr. Ser., vol 77, editedby [email protected]) M. S. Pringle,W. W. Sager,and W. V. Sliter,AGU, Washington,D.C. 1993. (Received June 6, 1994; revised November 11, 1994; Wolfe, C. J., M. K. McNutt, and R. S. Detrick, The Marquesas acceptedDecember 29, 1994.)