JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 94, NO. B8, PAGES 10,501-10,523, AUGUST 10, 1989

The Magellan Seamounts: Early Cretaceous Record of the South Pacific Isotopic and Thermal Anomaly

WALTER H. F. SM1TH, 1 HUBERT STAUDIGEL, 2 ANTHONY B. WATTS, 1 AND MALCOLM S. PRINGLE 3

We present geophysical, geochemical, and geochronological data from two Early Cretaceous seamounts in the westem Pacific Ocean which formed at melt sources in what is now French Polynesia. One seamount has an 40Ar/39Ar mineral age of 119.6 (+ 0.6, 1 •) Ma, is flexurally compensated with an elastic plate thickness Te of 10 (+ 2) km, and is isotopically similar to the island of Mangaia, or toward the HIMU manfie end-member. The second has an age of 100.1 (+ 0.5, 1 •) Ma, is compensated with a Te of 15 (+ 3) km, and is isotopically similar to Samoa, toward the EM-II "Dupal" mantle end- member. Geologic evidence suggests that the summits of these seamounts formed in shallow water, indicating that the underlying Jurassic seafloor was unusually shallow during the Cretaceous. Te values from these Cretaceous seamounts and localities in French Polynesia depart systematically from the simple square root of age law which otherwise applies throughout the ocean basins. A thermal rejuvenation of the lithosphere beneath these seamounts during or prior to their emplacement may explain their unusually low Te values and subsidence history. These two Early Cretaceous seamounts exhibit the isotope, Te, and depth anomalies which have been previously described at young volcanoes in French Polynesia, and we infer that a South Pacific Isotopic and Thermal Anomaly has existed for at least 120 m.y.

INTRODUCTION

Seamounts are the most abundant volcanoes on Earth, and one of the greatest concentrations of seamounts is in the western Pacific Ocean. Dredge samples and Deep Sea Drilling Project sites [Heezen et al., 1973] yield dominantly Cretaceous ages for these features. Studies of the seismically reverberant layer [Houtz and Ludwig, 1979], flexural loading [Watts et al., 1980], and seamount magnetic polarity [Hildebrand and Staudigel, 1986] suggest widespread magmatism in the Pacific during Cretaceous time. Studies of inferred hot spot traces [e.g., Clague and Jarrard, 1973; Epp, 1978; Duncan and Clague, 1985] and apparent polar wander paths [Harrison et al., 1975; Gordon, 1983; Hildebrand and Parker, 1987; Sager and Pringle, 1988] suggest that Cretaceous seamounts now located in the vicinity of the Magellan Seamounts may have originated in the region of French Polynesia.

The French Polynesia region includes the Marquesas, Society, Cook-Austral, and Pitcairn-Tuamotu seamount chains of the southcentral Pacific Ocean. Volcanic rocks

from these archipelagoes yield Neogene ages, generally less than 10 Ma (age data are summarized by Duncan and Clague [1985] and Calmant [1987]). Some volcanoes in the region are identified with a large-scale isotope anomaly (Dupal) in the underlying mantle [Duprd and Alldgre, 1983; Hart, 1984], while others indicate the presence of another component characterized by extremely radiogenic Pb

1Lamont-Doherty Geological Observatory of Columbia University, Palisades, New York, and Department of Geological Sciences, Columbia University, New York, New York.

2Scripps Institution of Oceanography, University of California, San Diego, La Jolla, California.

3Branch of Isotope Geology, U.S. Geological Survey, Menlo Park, California.

Copyright 1989 by the American Geophysical Union.

Paper number 89JB01080. 0148-0227/89/89JB-01080505.00

isotopes but low 87Sr/86Sr [Palacz and Saunders, 1986]. Although the Pb isotope systematics require long-term isolation of these magma sources (-109 years), they have been recognized only in young rocks, and the extent and location of these sources throughout geologic time is unknown. Calmant and Cazenave [1986, 1987] have shown that the volcanoes of French Polynesia are associated with anomalously low values of T e, the elastic thickness of the lithosphere, suggesting an unusual thermal structure in the lithosphere. Studies of seismic velocities [Nishimura and Forsyth, 1985; Dziewonski and Woodhouse, 1987], Seasat- derived gravity anomalies [Haxby and Weissel, 1986] and regional depth anomalies [McNutt and Fischer, 1987] indicate that this thermal anomaly may extend into the underlying as theno sphere. Like the geochemical evidence, the geophysical data reflect the present state of the lithosphere and asthenosphere beneath French Polynesia. If any of the anomalous features of French Polynesia are long- lived, we should expect to find evidence of them in Cretaceous volcanoes in the vicinity of the Magellan Seamounts.

In September 1985, R/V Conrad cruise 2610 surveyed a transoceanic cable route through the Magellan Seamounts and collected data from two seamounts near the cable route

(Figure 1). These two seamounts were selected because Watts et al. [1980] thought one to have formed "on-ridge" and the other "off-ridge." The more southerly seamount was named Hemlet Guyot by Smoot [1983], while the other is unnamed; we will refer to the latter as Himu Seamount, on the basis of its isotopic characteristics. The tectonic setting of these seamounts is shown in Figure 1. The crust is part of the Jurassic Quiet Zone to the south of the Japanese lineations that define the Pacific-Izanagi ridge crest during Late Jurassic time [Larson and Chase, 1972; Woods and Davies, 1982]. Handschumacher et al. [1988] have identified low-amplitude anomalies within the "Quiet Zone" in this area; Himu Seamount lies above their M33, while Hemlet Guyot is near their M36. Using the ages of M25 and M29 given by Kent and Gradstein [1985] and assuming a constant spreading rate prior to M25, we estimate seafloor ages of 162 and 165 Ma

10,501

10,502 SMITH ET AL.: MAGh7 J AN SEAMOUNT CRETACEOUS RECORD OF SOPITA

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Fig. 1. Index map for the study area in the western Pacific Ocean. Bathymetric contour interval 1000 m. Jurassic magnetic lineations and inferred fracture zones from Cande et al. [1978] and S. Cande (personal communication, 1985). The triangles and circles indicate seamounts identified as having formed "on-ridge" and "off-ridge", respectively, by Watts et al. [1980]. An AT&T cable survey passed between an on-ridge and off-ridge pair (Himu Seamount and Hemler Guyot, respectively) which are the object of this study.

under Himu Seamount and Hemler Guyot, respectively. In this paper we present geophysical data obtained during this cruise and a basic geochemical and petrological characterization of dredge samples, and we examine these data in light of the anomalous processes described at young volcanoes in French Polynesia.

DATA ACQUISITION AND REDUCTION

Navigation and Gravity

Navigation data were digitally recorded once per minute from Transit satellite, Global Positioning System (GPS), and LORAN receivers and from pit, gyrocompass, and Doppler

speed logs. The Transit displays a position dead reckoned since the last satellite fix; fixes come roughly once per hour. LORAN receivers were operated in range-range mode with a Rb vapor frequency standard for reference. The range calibration to transmitters was adjusted to GPS fixes when necessary. Reliable GPS data were available only a few hours per day, and the ship surveys were navigated primarily with LORAN and Transit positions.

Gravity data were obtained with a Bell-Aerospace BGM-3 sea gravimeter mounted on a gyrostabilized platform. The axisymmetric design of the sensor eliminates cross-coupling errors. An evaluation of the precision, accuracy, and capabilities of this instrument has been made by Bell and Watts [1986], who showed that the accuracy of the system is limited by the quality of navigation, and that the instrument

SMITH E'r AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,503

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Fig. 2. Map of the ship track over the target seamounts and histograms of crossover errors in free-air gravity anomalies derived from three navigation solutions: LORAN, DR, and GPS. Gravity anomalies along profiles A-A' and B-B' appear in Figures 8 and 12.

is capable of submilliGal accuracy (1 mGal = 10 -5 m s -2) at wavelengths of a few kilometers when adequate navigation data are available.

After the cruise, we analyzed three navigation solutions: GPS data, LORAN positions corrected for a linear drift between calibrations, and a third solution which we refer to

as DR. DR is essentially a dead reckoning scheme in which the Doppler speed log is used to reckon between representative fixes selected from visual analysis of scatter plots of the GPS, LORAN, and Transit data files. Whereas in the past, dead reckoning usually spanned hours between celestial or satellite fixes, this DR is supplied fixes approximately every 15 min.

Our preferred technique to assess navigation quality is to examine the crossover error (COE) between observed gravity anomaly values at intersecting ship tracks within a survey [Wessel and Watts, 1988; Wessel, 1989]. Although COEs in any geophysical signal might be used, we feel that gravity

COEs are the best test because the calculated gravity anomaly is linked to the navigation through the E6tv6s effect. The recorded raw gravity signal contains both the actual gravity field and the E6tv6s effect arising from the ship's motion with respect to the rotating earth [Dehlinger, 1978]. An E6tv6s correction to the raw gravity must be calculated from the navigation data. This correction is a function of the ship's velocity (heading and speed) over the ground and thus requires the time derivative of the navigation which amplifies noise in this series. Over a seamount, the gravity gradients may reach several milliGal per kilometer, and so a small COE in gravity indicates that both the navigation series and its derivative velocity are accurate. Low COEs in other geophysical data would show only the accuracy of point positions at crossovers but not the quality of the navigation time series as a whole.

Figure 2 shows the survey track and histograms of the COEs in gravity obtained by each of the three navigation

10,504 SM1TH E• AL.: MAGELLAN SEAMOLrNT CRETACEOUS RECORD OF SOPITA

EStvSs effect

mGal

Uncorrected gravity a

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Reduced qravity

100

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4 hr 4 hr Y

Sept. 18,1985 R/V Conrad 2610 Sept. 18,1985 R/V Conrad 2610

Fig. 3. Time series plot of the recorded raw gravity and the E6tv6s effect and resulting free-air anomaly calculated for the three navigation solutions of Figure 2. Note the nature and level of the errors in each signal. (a) E6tv6s effects: At X the LORAN fails to follow a series of course adjustments. At Y the GPS makes an "excursion." (b) Free-air anomalies: At X the LORAN error produces a spurious anomaly and tears of 25 mGal. At Y the GPS excursion produces erroneous anomalies up to 50 mGal.

schemes. The number of crossovers found in each case is

not the same for two reasons. (1) The GPS data exist for only a few hours each day and only crossovers on these path segments are counted. These show the highest rms (root mean square) gravity COE, however. (2) The LORAN navigation contains a large number of crossovers because there is a large scatter in the positions and the track appears to cross itself many times during dredging operations. Excluding these crossovers does not improve the rms COE of 7 mGal, however. The DR solution gives the best rms COE, less than 2 mGal, and does not have difficulties during dredging operations.

Figure 3 illustrates how navigation errors enter the free-air gravity anomaly through the E6tv6s correction. Generally, sharp changes in raw gravity indicate changes of the ship's course and should be mirrored in the E6tv6s signal. At X a series of course adjustments begins, resulting in the climb in observed gravity. The DR and GPS E6tv6s signals reflect this (Figure 3a), but the LORAN fails to do so. At Y the GPS data show the ship making an "excursion" which the DR and LORAN signals do not indicate. The raw gravity shows that the ship did not make this excursion. Figure 3b shows the free-air anomalies derived from the E6tv6s signals of Figure 3a. The noise from the LORAN signal is evident, as is its failure to correct for ship motion at X. The GPS excursion at Y produces errors as large as 50 mGal in the anomaly. Such excursions were common throughout our survey and account for the very poor COEs derived with GPS data. We therefore used the DR navigation for all subsequent data analysis.

Sea Beam Bathymetry

A vertical reference was supplied to the Sea Beam system from the gravity meter gyrostabilized platform. Expendable bathythermographs (XBTs) were used to compute a sound velocity profile in the water column and correct the slant range of off-vertical beam signals for path refraction. Further details of the Sea Beam system are discussed by Renard and Allenou [1979]. Our surveys were laid out to collect gravity data along traverses of the flexural moats surrounding the target seamounts, to dredge the flanks of the targets, and to obtain good Sea Beam coverage of Himu Seamount. A multibeam survey of Hemler Guyot was previously made by Smoot [1983]. We did not attempt to duplicate his results.

Nominal depths reported in this study are derived from travel time using a uniform velocity of 1.5 km s-1. After the navigation reduction and analysis, the DR navigation solution was merged with the Sea Beam data for final plotting. Maps produced at a scale of 1'180,000 and contoured at 20 m show no significant bathymetric errors at overlapping swaths.

Dredge Sampling

We placed six dredges, two on Himu Seamount and four on Hemlet Guyot. Dredges were navigated with Sea Beam, and locations, depths, and amount recovered are shown in Figures 4 and 5 and listed in Table 1. All dredges generally contained fragments heavily encrusted with manganese oxide. Pillow basalts and flows were the dominant rock

SMITH E-T AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,505

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Fig. 4. Bathymetry at Himu Seamount from Sea Beam data collected in this study. Contour interval 100 m; solid contours are from Sea Beam collected along finely dotted ship tracks; dashed contours are inferred with reference to wide beam echo soundings collected by Verna cruise 33-12 (coarse dotted lines). Dredge locations indicated by circles.

types recovered from Himu Seamount, while submarine volcaniclastics and carbonates were frequently recovered from Hemler Guyot. Dredges were placed on Himu Seamount on the south (RD 3) and north (RD 4) flanks. Dredges were obtained from Hemlet Guyot on the southwest rift zone of the pyramidal peak on the south flank of the main guyot (RD 5), on the northeast rift zone of the main guyot (RD 6), on the southwest flank of the main guyot (RD 7), and on the northwest flank of the northern guyot in the complex (RD 8).

MORPHOLOGY

A 100-m contour map of Himu Seamount is shown in Figure 4. The 5600-m contour approximately defines the base of the seamount. This contour is ovate in shape, 45 by 60 km, with a long axis that strikes approximately N35øW. The seamount has two peaks at 2000 m depth which exhibit distinct radiating fissures, or rift zones [cf. MacDonald, 1972], on their flanks. The southern summit has four prominent rift zones which strike N45øE and N35øW, giving

10,506 SMITH •l' AL.: MAG• SEAMOUNT •ACF_.DUS RECORD OF SOPITA

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Fig. 5. Bathymetry of Hemler Guyot after N. C. Smoot (personal communication, 1985). Contour interval 100 fm (183 m). Our survey tracks are shown as dotted lines. Dredge locations indicated by circles.

it a rectangular pyramidal shape. The saddle point between the summits lies at 3480 m depth, and slopes from the summits to the saddle point are as steep as 17'; on the seamount flanks there are breaks in slope at 4700 and 5000 m, and lower slopes are typically 4*-5 ø.

Figure 5 is a map of Hemlet Guyot contoured at 100 fathoms (100 fm = 183 m) and based on multibeam bathymetric data held by NAVOCEANO (N.C. Smoot, personal communication, 1985). The feature has three summits' two flat topped guyors with a shelf break at approximately 1400 m depth and a rhombic pyramidal peak at 1500 m. For comparison, this last is somewhat larger in size than the peaks of Himu Seamount. The largest edifice is an elongate guyot 35 by 10 km at the top, with a long axis that strikes N45øE. A circular guyot, 8 km in diameter, is

19 km north of the main guyot. Both guyots have a complex pattern of radiating rift zones on the flanks dominated by a N40ø-50øE trend. The rhombic peak is 45 km SE of the principal edifice; its rift zones strike N45øE and N55øW and slopes exceed 22 ø. Saddles between summits lie at 3400 m depth in the north and 4300 m in the southeast. Slopes break at 3400 and 4500 m on the flanks. The depth to the seafloor is approximately 5600 m, and the base of the complex is 130 km N-S by 110 km E-W.

Since our ship time was limited to 5 days and a multibeam survey of Hemlet Guyot had been previously carried out by Smoot [1983], we concentrated our survey time on Himu Seamount. We made only a few crossings of Hemler Guyot to collect gravity data and dredges. Sea Beam swaths on these crossings agree well with Smoot's map except at the

SMITH ET AL.: MAG• J AN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,507

TABLE 1. Dredge Locations, Depths, and Approximate Amount of Material Recovered During Cruise RC2610

Depth, m Approximate Approximate Amount Dredge Start End Latitude Longitude Recovered, kg

Himu Seamount

RC26 D3 4520 3400 N21ø28 ' E151ø47 ' 200

RC26 D4 4500 3440 N21ø42 ' E151ø42 ' 80

Hernler Guyot

RC26 D5 4480 3540 N19ø22 ' E151ø58 ' 100

RC26 D6 4220 4100 N19ø53 ' E151ø54 ' 30

RC26 D7 3950 2510 N19ø33 ' E151ø34 ' 150

RC26 D8 4260 3400 N20ø06 ' E151ø34 ' 250

circular satellite guyot on the north flank of Hemlet Guyot. We manually adjusted the bathymetric contours at this flank guyot to match our reconnaissance swaths and found that this adjustment also improved the fit of our gravity models. We believe Figure 5 is an accurate representation of Hemlet Guyot; however, NAVOCEANO has classified the raw data used in Smoot's map, and the quality and coverage of the data control for his map cannot be ascertained.

PE'rROGRAPHY AND GEOCHEi•STRY

Himu Seamount

The dredges from Himu Seamount (Table 1) yielded samples of angular and subangular basaltic flows with one basaltic lapillistone, all typically encrusted with 1-4 cm thick Mn crusts. Dredge 3 contained only aphyric or slightly clinopyroxene/amphibole porphyritic (<2% vol) samples. Vesicle contents are low (0.5%) with commonly developed segregation vesicles. Samples from dredge 4 were petrographically more variable, ranging from aphyric basalts (RC26 4-5) to extremely porphyritic amphibole- clinopyroxene basalt (25% amphibole, 5% clinopyroxene; RC26 4-8). Moderately porphyritic types include a trachybasalt with 3% plagioclase, 2% amphibole, and 1% nepheline phenocrysts (RC26 4-1). Most extrusive rocks contain less than 10% vesicles, rarely up to 20%; segregation vesicles are common.

The most common phenocryst phase found in dredges 3 and 4 is a kaersutitic amphibole, followed by diopsidic/salitic clinopyroxene and labradorite plagioclase (see electron microprobe analyses in Table 2). Diopsidic to salitic clinopyroxenes display very good correlations in A1 versus Si and Ti versus Si diagrams, covering a large range as is typical for alkali basaltic suites from oceanic intraplate volcanoes (Figure 6). No olivine phenocrysts nor their altered remains were observed. Groundmass textures range

from intersertal to microcrystalline matrix with plagioclase, amphibole, and submicroscopic mesostasis. All samples are very altered, particularly the aphyric, fine-grained, and more vesicular samples. However, clinopyroxene and amphibole are mostly unaltered, and more than 50% of the feldspar is

preserved. Secondary mineral assemblages are dominated by smectites but also include opaques and zeolites. The samples from Himu Seamount represent a suite of alkalic basalts, as evidenced by the common titaniferous amphibole and clinopyroxene and the (rarer) occurrence of nepheline phenocrysts or trachytic groundmass textures.

Hemler Guyot

At Hemler Guyot, three major peaks were dredged (Table 1). Dredge 5 recovered heavily (up to 6 cm thick) Mn- encrusted angular pillow basalt fragments, interpillow 1.,,,,,1 .-, .-•1 .,,,, ' ,,,• ...... tlte, and incipient pillow breccia (nomenclature as by Fisher and Schrnincke [1984]). The basalts are mostly aphyric, occasionally with 0.3 mm microphenocrysts of clinopyroxene (up to 10% Ti augite) and up to 0.05 mm euhedral olivine (< 2%) that is completely replaced by a reddish material similar to iddingsite. Electron microprobe analyses suggest that these Hemlet Guyot clinopyroxenes contain less Ti than clinopyroxenes from Himu Seamount (Table 2). Groundmass textures range from intersertal to subophitic with 5-20% vesicles. All samples from this dredge are very altered.

Dredge 6 from the northeast flank of the main guyot recovered only two big blocks, a sideromelane lapillistone and one basalt fragment with 5-10% olivine and 5% plagioclase microphenocrysts. Both samples are very altered.

Dredge 7 from the south flank of the main guyot included a large assemblage of volcaniclastic and extrusive rocks, all variably encrusted with manganese. Volcaniclastics include well-sorted, monolithological hyaline lapillistones, fine- grained hyaloclastites, pillow fragment breccias, and lapillistones cemented by limestone or Mn oxides. The extrusives are all alkali basalts with phenocrysts of clinopyroxene (often Ti augite with well-developed anomalous interference colors), olivine (100% altered), and minor plagioclase, and rare potassium feldspar xenoliths (RC26 7-8; see Table 2 for typical mineral analyses). The rock types range from ankaramites with up to 25% Ti augires and 15% olivine (100% altered; RC26 7-16) to hawaiitic [RC26 7-8) alkali basalt types. Up to 15% of (completely replaced) olivine phenocrysts, minor plagioclase, and alkali

10,508 SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA

TABLE 2a. Microprobe Analysis of Amphibole Phenocrysts from Alkalic Basalts from Himu Seamount and Hemler Guyot

Sample SiO 2 TiO 2 A1203 FeO MgO MnO CaO Na20 K2 ̧ Cr • Sum

RC26 4-1 39.64 5.68 13.17 12.2 11.54 0.19 11.92 2.65 0.92 0.01 2.02 99.95 RC26 4-6 39.54 5.64 13.55 11.52 12.07 0.18 11.76 2.72 0.94 0.01 2.02 99.94 RC26 4-8 39.59 5.03 13.33 12.9 11.5 0.19 11.69 2.68 0.96 0.01 2.01 99.85

Sample Si A1 Ti Fe Cr Mg Mn Na Ca K • Sum

RC26 4-1 5.896 2.31 0.636 1.518 0.001 2.558 0.024 0.765 1.9 0.174 1. 16.781 RC26 4-6 5.862 2.367 0.628 1.429 0.001 2.667 0.023 0.782 1.868 0.177 1. 16.804 RC26 4-8 5.908 2.344 0.564 1.61 0.001 2.556 0.024 0.775 1.869 0.182 1. 16.833

TABLE 2b. Microprobe Analysis of Clinopyroxene Phenocrysts from Alkalic Basalts from Himu Seamount and Hemler Guyot

Sample SiO 2 TiO 2 A1203 FeO MgO MnO CaO Na20 K20 Cr Sum

RC26 4-8 49.19 1.415 4.815 8.338 12.56 0.205 22.19 0.633 0.003 0.118 99.46 RC26 4-8 47.15 2.485 6.935 6.87 12.73 0.125 22.57 0.55 0 0.045 99.45 RC26 4-6 46.25 2.53 8.6 7.24 12.45 0.125 21.4 0.635 0.01 0.17 99.4 RC26 7-8 48.26 1.917 6.357 5.88 13.67 0.087 22.42 0.527 0.017 0.417 99.54 RC26 8-8 51.57 0.917 2.963 6.01 16.63 0.113 20.92 0.257 0.013 0.35 99.75

Sample Si A1Z A1 Fe Mg Mn Cr Ti Na K Ca Sum

RC26 4-8 1.852 0.148 0.065 0.263 0.704 0.007 0.004 0.04 0.046 2E-04 0.895 4.0229 RC26 4-8 1.769 0.231 0.076 0.216 0.712 0.004 0.001 0.07 0.04 0 0.907 4.0261 RC26 4-6 1.735 0.265 0.115 0.227 0.696 0.004 0.005 0.071 0.046 3E-04 0.86 4.0242 RC26 7-8 1.798 0.202 0.077 0.183 0.759 0.003 0.012 0.054 0.038 8E-04 0.895 4.0218 RC26 8-8 1.901 0.099 0.03 0.185 0.914 0.004 0.01 0.025 0.018 6E-04 0.826 4.0133

0.2771

0.12767

0.24955 0.19753 0.17117

Sample Wo En Fs

RC26 4-8 0.479 0.377 0.144 RC26 4-8 0.494 0.387 0.119 RC26 4-6 0.481 0.389 0.129 RC26 7-8 0.486 0.413 0.101 RC26 8-8 0.428 0.474 0.098

Read 2E-04 as 2 x 10 -4.

TABLE 2c. Microprobe Analysis of Plagioclase Phenocrysts from Alkalic Basalts from Himu Seamount and Hemler Guyot

Sample SiO 2 A1203 FeO MnO MgO CaO Na20 K20 Sum

RC26 4-1 54.33 28.91 0.468 0.008 0.048 11.05 4.827 0.415 100.1 RC26 7-8 66.99 19.62 0.117 0.013 0.003 0.147 7.257 6.077 100.2 RC26 8-8 53.97 28.65 0.638 0.006 0.101 11.46 4.611 0.414 99.86

Sample Si A1 Fe Mg Na Ca K Mn Sum ,

Ab Or

RC26 4-1 2.453 1.539 0.016 0.003 0.423 0.535 0.024 3E-04 4.992 43.05 54.513 2.4332 RC26 7-8 2.979 1.028 0.004 3E-04 0.626 0.007 0.345 5E-04 4.989 64.04 0.7204 35.244 RC26 8-8 2.455 1.53 0.019 0.005 0.412 0.543 0.022 3E-04 4.987 42.14 55.584 2.2726

Read 2E-04 as 2 x 10 '4.

SM1TH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,509

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Himu Seamount rims

Hemler Guyot cores

Hemler Guyot rims Ocean Island Alkali Basalt

Ocean Island Tholeiite

Si

Fig. 6. Si versus Ti per formula unit for clinopyroxenes from electron microprobe analyses of phenocrysts in samples from Hemlet Guyot and Himu Seamount in reference to clinopyroxene data from ocean island and seamount alkali basalts and tholelites after Fodor et al. [1975] and Schweitzer et al. [1979].

feldspar xenocrysts also occur (RC26 7-8). Vesicle contents of extrusives are typically between 5 and 20%.

Dredge 8, from the small satellite guyot on the north flank of Hemlet, also recovered a variety of basalt fragments and volcaniclastics, the latter including lapillistones, pillow fragment breccias, and other clastic rocks containing limestones and volcaniclastics. Basalts include

clinopyroxene and olivine and clinopyroxene and plagioclase as primary phenocryst assemblages. The groundmass is generally very fine grained and extremely altered. Vesicle contents are typically between 5 and 10%.

All lithologies recovered from both seamounts were formed in a submarine environment, as evidenced by pillow structures, lack of well-rounded fragments, and quench textures. Dredges from guyors recovered more volcaniclastics, recording shallower eruption depths. Based on petrography and mineral compositions, all samples recovered are characteristic for normal alkalic ocean island

basalts; the samples from Himu Seamount appear to be more alkalic than the samples from Hemlet Guyot, with a very large overlap. By analogy to Hawaiian basalts, most rocks dredged may reflect the late shield building or posterosional alkalic stage of volcanism. One sample with large nepheline phenocrysts may be part of a late nephelinitic stage, as described in Hawaii by Clague and Dalrymple [1987].

Isotopic Analysis

We have analyzed some representative samples for Sr, Nd, and Pb isotope ratios, using techniques as in the work by Zindter et al. [1984] (Table 3). All samples were extensively leached with large acid/rock ratios to remove sea

floor alteration effects. We show our data in a Pb-Sr isotope variation diagram with respect to other oceanic basalt fields in Figure 7. The shaded fields in Figure 7 are from French Polynesian and Samoan samples [Vidal et al., 1984; White, 1985; Palacz and Saunders, 1986; Dupuy et al., 1987; Wright and White, 1987] which are interpreted to delineate mixing between the EM-II and HIMU mantle end-members [Zindler and Hart, 1986]. Also shown in Figure 7 are analyses of samples from Erikub and Ratak Guyots located in the Marshall Islands to the southeast of the Magellan Seamounts (after Staudigel et al., [1987, The Longevity of the South Pacific Isotopic and Thermal Anomaly, submitted to Earth and Planetary Science Letters, 1989, hereafter referred to as 1989 submitted] and Davis et al. [1989]). The data from Hemlet Guyot lie toward the EM-II end-member and are similar to samples from Samoa, the Marquesas, and Aitutaki (Cook Islands); the Himu Seamount analyses plot toward the HIMU end-member in the same field as samples from Rurutu (Austral Islands).

The EM-II end-member (enriched mantle, type II) is characterized by low 143Nd/144Nd, high 87Sr/86Sr, and high 207pb/204pb and 208pb/204pb at a given value of 206pb/204pb and is thought to be derived from injection into the manfie of continental sediment or crust or oceanic

island material in subduction zones (for review, see Zindler and Hart [1986]). While the most extreme EM-II hot spots lie in the southern hemisphere (the "Dupal anomaly" [Duprd and Alltigre, 1983; Hart, 1984]), the component has also been described in other localities, for example, at Shimada Seamount [Graham et al., 1988]. Anderson [1982] postulated Pangaeatic subduction for the source of the Dupal anomaly. The HIMU (high l• = 238U/204pb) end-member is

10,510 SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA

TABLE 3. Isotope Ratios of Selected Samples

Sample 87Sr/86Sr 143Nd/144Nd 206pb/204pb 207 pb/204pb 208 pb/204pb

Himu Seamount

RC26 3-1 0.702949 0.512894 20.574 15.764 39.728

RC26 4-1 0.702782 0.512858 20.671 15.764 39.94

Hemler Guyot

RC26 5-5 0.704686 0.512685 19.29 15.631 39.313

RC26 6-1 0.705065 0.512722 19.017 15.636 39.213

RC26 7-8 0.704761 0.512758 19.294 15.646 40.592

RC26 8-8 0.707718 0.512759 18.907 15.664 39.256

Accuracy of isotope ratios: Pb, 0.1% relative; Sr, 0.000035 absolute; Nd, 0.000025 absolute.

characterized by high 206pb/204pb, 207pb/204pb and 208pb/204pb at relatively low 87Sr/86Sr, suggesting enrichment in U + Th relative to Pb (but not Rb relative to Sr) of the order of 109 years ago; subduction of ancient altered oceanic crust has been suggested as the source of this

material (for review, see Zindler and Hart [1986]). The mixing between these two end-members and their close spatial association in the Cook-Austral island groups suggest that they may share a common origin. Trace element analyses of Cook-Austral basalts suggest that their

17 18

0.708 I

0.707 -

0.706 -

0.705 -

0.704 -

0.703 -

0.702

[7

206 Pb / 204 Pb 19 20

I I

ß I EM-II

21 22 I

i i i i

HIMU

18 19 20 21 22

206 Pb / 204 Pb

0.708

0.707

0.706

0.705

0.704

0.703

0.702

Fig. 7. Sr-Pb isotopic variation in the Magellan seamounts compared with oceanic basalt samples. Shaded fields are from south Pacific islands [Vidal et al., 1984; White, 1985; Palacz and Saunders, 1986; Dupuy et al., 1987; Wright and White, 1987]. Light grey, Samoa; medium grey, Marquesas; dark grey, Cook-Austral samples. Circles are from Hemler Guyot, and diamonds are from Hirnu Seamount. Squares are from samples in the Marshall Islands concordant with the hot spot track from Himu Seamount [Staudigel et al., 1987; Davis et al., 1989].

SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,511

source has previously undergone extraction of island arc tholelite [Palacz and Saunders, 1986; Dupuy et al., 1987; Saunders et al., 1988]. This is consistent with the suggestion of Hart and Staudigel [1989] that altered oceanic crust cannot generate a HIMU mantle end-member without compositional modification during the subduction process. This lends further support to the idea that present-day French Polynesia may overlie mantle modified at an ancient subduction zone. The above inferences are based largely on samples from young volcanoes in French Polynesia. An important result of our data is that these same mantle signatures have been found in Cretaceous basalts, and that in our samples mixtures near the end-members EM-II and HIMU are again found in close association.

GEOCHRONOLOGY

The K-Ar method has been widely used to date oceanic lavas, but the samples dredged from Himu Seamount and Hemler Guyot are too altered to yield reliable whole rock ages or sufficient material to separate minerals for conventional ages. However, we were able to separate

sufficient quantities of phenocryst phases from some samples that, with careful acid cleaning, proved suitable for ;40Ar/39Ar total fusion analysis. Hornblende was recovered in sufficient quantities (0.4-0.8 g) to allow analysis following conventional induction heater fusion after the techniques described by Dalrymple and Lanphere [1971]. Plagioclase and nepheline were recovered in sufficient quantities (0.0001-0.002 g) to allow analysis by laser fusion with the Great Little Machine (GLM) following the techniques described by Dalrymple and Duffield [1988]. The ability to analyze samples as small as 0.1 mg has only recently become available, and it enables us (1) to determine ages for samples previously limited by insufficient sample quantity, and (2) to achieve better precision than was previously possible even with much larger samples.

To establish a reliable age for a seamount, we feel it is critical to examine at least two different samples, representing different lithologies wherever possible. Thus we have determined ages on five phenocryst phases from four different lithologies, two each from Himu Seamount and Hemlet Guyot (Table 4). The best age for each seamount is the weighted average of the individual analyses, where the

TABLE 4. 40Ar/39Ar Total Fusion Ages of Mineral Separates From the Magellan Seamounts

Sample Material J 40Ar/39Ar 37Ar/39Ara 36Ar/39Ar 36Arca b 39Arca b Ma + 1 s.d.

RC26 4-1 hornblende 125-500

Himu Seamount

Weighted age

0.002656 30.653 5.7717 0.01718 8.87 0.39 0.02 84.9 121.0 + 0.8

plagioclase 125-250 gm 0.00449

RC26 4-8 hornblende 250-500 [tm 0.002656 250-500 lain 0.002656

RC26 7-5/8 nepheline 125-250 IJxn 0.005137

/

Weighted age

125-250 gm 0.00448 RC26 7-15 plagioclase

Weighted age

15.723 12.843 0.00483 66.76 0.86 0.24 96.7 120.2 q- 0.8

15.639 12.797 0.00438 73.31 0.86 0.24 97.5 120.5 q- 2.6 15.724 13.015 0.00487 67.06 0.87 0.24 96.7 120.2 q- 1.0 15.745 14.516 0.00665 54.77 0.97 0.24 94.1 117.3 + 1.0 15.624 14.089 0.00472 74.86 0.95 0.24 97.5 120.5 q- 0.9 15.964 14.273 0.00685 52.27 0.96 0.24 93.7 118.4 q- 1.3

119.6 q- 0.7

31.025 5.7963 0.01973 7.76 0.39 0.02 82.7 119.3 q- 0.8

30.650 5.7027 0.01823 8.26 0.38 0.02 83.9 119.6 q- 0.8

Hemler Guyot

11.020 0.02535 0.0001819 3.70 0.002 0.05 99.5 98.8 q- 0.7 11.209 0.02454 0.0002349 2.76 0.002 0.05 99.3 100.4 q- 0.7 11.115 0.01481 0.0000133 29.3 0.001 0.05 99.9 100.1 q- 0.8 11.081 0.03387 0.0001626 5.50 0.002 0.05 99.5 99.4 + 0.8 11.004 0.00788 0.0000198 10.5 0.001 0.05 99.9 99.1 q- 0.9 11.169 0.02065 0.0004696 1.16 0.001 0.05 99.7 99.4 q- 1.0

11.093 0.07488 0.0001992 9.93 0.0{}5 0.05 99.5 99.5 q- 1.3 11.151 0.09095 0.0000370 64.5 0.006 0.05 99.9 100.4 q- 1.5

99.5 + 0.6

13.332 1.5626 0.002557 15.34 0.10 0.29 94.9 99.6 + 0.9 13.331 1.083! 0.001255 21.66 0.07 0.29 97.5 102.2 q- 1.3 3.298 0.7551 0.001572 12.06 0.05 0.29 96.6 101.0 q- 0.9

100.7 q- 0.6

aCorrected for 37Ar decay, half-life = 35.1 days. b b Su scripts indicate radiogenic (R), calcium-de, r•ved ,(Ca), and potassiurr!,•deriv,ed (K) argon, respectively. c io i IO 1 Ages were calculated using •,• = 0.581 x 10' yr- , •,[• = 4.962 x 10- yr- , 40K/K-total = 1.167 x 10 '4 mol/mol; errors are estimates

of the standard deviation of analytical precision with a long-term limiting precision of 0.5%. Weighted ages are weighted by the inverse of the variance.

10,512 SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA

weighting factor is the reciprocal of the variance [Taylor, 1982], and the limiting precision is based on a long-term estimate of analytical precision of 0.5% (1 s.d.).

For Himu Seamount, the best age is the weighted average of the three individual hornblende analyses and the average laser fusion plagioclase analysis: 119.6 + 0.6 Ma. For Hemler Guyot, the best age is the weighted average of the two laser fusion average ages: 100.1 + 0.5 Ma. These uncertainties are the one standard deviation analytical precision; if these seamounts fit the Hawaiian growth model, in which an individual volcano is built in less than 0.5 m.y. and an entire edifice is built in about 1 m.y. [Clague and Dairytopic, 1987], then the 95% confidence interval for these age estimates would include the entire volcanic history of each of these seamounts. On the other hand, if the rocks

we have dated are from a posterosional or alkalic-rejuvenated stage [Clague and Dairytopic, 1987; 1988], then the bulk of each seamount could be 3-5 m.y. older than the given dates. Finally, we note that age relationships among volcanoes in the Cook-Austral island chain are more complex than the Hawaiian model; ages of rocks from Rarotonga Island, for example, span about 12 m.y. [Turner and Jarrard, 1982; Duncan and Clague, 1985].

ISOSTATIC COMPENSATION

Methodology

The response function technique [Watts, 1978] provides an efficient means to compute the gravity effect of a seamount and its flexural compensation. For a three dimensional solution this requires gridded data sets. The grid spacing must be smaller than the shallowest bathymetry value for the discrete Fourier transform approximation of potential anomalies to hold [Parker, 1972]; for our seamounts this means a grid spacing less than 1.4 km. At the same time, in order that all flexural wavelengths be included, the grid should extend beyond the feature of interest at least 400 km. Thus a complete solution by this method requires large arrays. Survey coverage is rarely complete, and the depth values at unsurveyed grid locations must be interpolated; as the grid spacing decreases, many more interpolated points are required. The gridding process may degrade the rich detail of the Sea Beam bathymetry.

Line integral methods [Talwani and Ewing, 1960] may be used to compute the gravity effect at any point of three- dimensional bodies of arbitrary shape. The body is described by topographic contours which are approximated by polygons. The gravity effects of thin laminae bounded by polygons are computed and then integrated over depth to produce the calculated anomaly. The accuracy of this method is limited only by the number and detail of the polygon contours available.

We have combined these two techniques, using the lamina method to compute the gravity effect of the seamount and the Fourier method to compute the gravity effect of flexural compensating surfaces under the seamount. This combination has several advantages. First, it makes full use of the accuracy and detail of the Sea Beam data. Second, it allows geological interpretation as the user chooses how to close contours where survey coverage is incomplete. Third, gridded data are used only to compute gravity effects from

flexed surfaces in the crust, and since the short-wavelength topography of these surfaces is damped by flexure, the grid spacing may be increased to the minimum seafloor depth (greater than 5 km), reducing the size of the required arrays and the number of empty grid points which must be interpolated. Fourth, although the gravity effect of the topography is computed from a large number of detailed contours, it is linear in density and does not vary with flexural rigidity so it need only be computed once.

We digitized the contours in Figures 4 and 5 and used these points to fill in gaps in the Sea Beam swath coverage. These combined data were then gridded at 2.5 arc min using a minimum curvature algorithm [Briggs, 1974; Swain, 1976]. We also tried using the 5 x 5 min SYNBAPS bathymetry [Van Wykhouse, 1973] to fill in gaps but found this unsatisfactory because the target seamounts were mislocated and poorly represented in the SYNBAPS data.

Determination of T e

We computed forward models of the gravity fields of the seamounts using three values of load density (p = 2.4, 2.6, and 2.8 g cm-3) and seven values of Te, the elastic thickness of the lithosphere (0, 2, 5, 10, 15, 20, and 25 km). The T e = 0 corresponds to Airy isostasy. Additional parameters used are summarized in Table 5. We evaluated each model by visual analysis of profiles along track and quantitative analysis of misfit statistics, as discussed below.

Himu Seamount. Of the models computed for Himu Seamount, T e = 10 km and load density = 2.6 gcm -3 best fits the observed gravity anomaly, according to our visual analysis of the data. This model is shown in Figure 8 along profile A-A' of Figure 2. In order to place confidence limits on T e we investigated quantitative measures of the goodness of fit of the models. The root-mean-square (rms) amplitude of a model's gravity residual is a commonly used goodness of fit criterion. The rms for all models is shown in Figure 9a, contoured at 0.5 mGal. A broad region of the Te-density parameter space has an rms less than 5 mGal, and several values of T e yield an rms between 3.0 and 3.5 mGal, including the choice we made visually. (Note that this misfit value is approximately twice the rms crossover error.) The rms misfit can be used to rule out the worst models, but it is a poor discriminant of T e, because it does not contain

TABLE 5. Summary of Parameters Used in Gravity Model Calculations

Parameter Value

Load density 2.4, 2.6, or 2.8 g cm-

Infill density = load density

Mantle density 3.33

Layer 2 density 2.8

Layer 3 density 2.9

Layer 2 thickness 1.5 km

Layer 3 thickness 4.5 km

Young's modulus 1011 Pa Poisson's ratio 0.25

SMITH Er AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,513

50-

Gravity residual

0 •

-50 -

1 O0 -

50-

0 i

100 -

50

0-

2000 -

4000 -

6000 -

Topography

I I I I I 0 25 50 75 100 1•5

density 2.6 T - 10 km

e

observed

Fig. 8. Profile along A•A' of Figure 2 showing the fit of the gravity model at Himu Seamount.

information about the shape, or waveform, of the residual anomaly.

A measure of fit which is sensitive to shape is the correlation of the residual gravity with the topography, shown in Figure 9b contoured at 10% intervals. Models with density too high predict a gravity anomaly too large, and so the residual is negative and negatively correlated with the topography; the converse is true where the density is too low. Overcompensation or undercompensation has an analogous effect. Because the correlation changes sign, there is a zero contour in Figure 9b; however, no model produces a perfect (zero) correlation. Of the models computed, the one closest to zero correlation is also the one we chose visually.

We have combined the rms and correlation coefficient fit measures into an "objective function" as follows:

obj. fn. = 0.5 , [rms/rmSmax+abs (corr. coeff.)]

where rmsma x is the maximum rms of all the models; we use this to normalize the rms range from 0 to 1. We take the absolute value of the topography correlation coefficient so this ranges from 0 to 1. The objective function then ranges from 0 to 1. This is contoured in 10% intervals in Figure 9c, and illustrates that the only model which minimizes both rms and topography correlation is the one that we chose on visual grounds. Our experience modelling T e leads us to believe that T e is between 8 and 12 km at this

10,514 SMITH ET AL.: MAGEI JAN SEAMOUNT CRETACEOUS RECORD OF SOPITA

seamount, or within the 15% contour of the objective function in Figure 9c.

Figure 10 is a plan view of the residual anomaly from the best fitting model. There is a small -4-mGal low centered over the northern summit and a broad +8-mGal high centered over the southeastern flank. Studies of an uplifted seamount, La Palma [Staudigel and Schmincke, 1984], suggest that intrusive rocks make up more than 90% of the deep structure of seamounts and that the shallow structure may consist chiefly of volcaniclastic rocks if the summit is erupted in shallow water (approximately 800 m or less). Hill and Zucca [1987] have inferred a dike complex in the flank of Hawaii from seismic and gravity data. We feel that the residual gravity field at Himu Seamount indicates the presence of

~9..0

such features. Densities appropriate for hyaloclastites S) in gcm -3) in the summit areas and gabbros (-3.0 g cm- the southeast flank can fit the gravity residual in Figure 10.

Hemler Guyot. The objective function for Hemler Guyot is shown in Figure 11. Of the suite of models calculated, T e = 15 km and load density of 2.8 gcm -3 gives the best fit. However, unlike the Himu Seamount objective function (Figure 9c), this one shows no definitive minimum. We did not try load densities exceeding 2.8 gcm -3 since we do not expect the average bulk densities of seamounts to exceed this value. In modelling the gravity field of Hemlet Guyot

2.8

2.6

2.4

0 5 10 15 20 25

2.8

2.6

2.4

2.8

2.6

2.4

I / I I

_

0 5 10 15 20 25

Te (km)

Topography correlation

I I I I I

0 5 10 15 20 25

Te (km)

Te (km)

RMS residual amplitude

Fig. 9. Quantitative measures of gravity model misfit for Himu Seamount. (a) The rms of residual gravity anomaly. A broad region of the model parameter space produces an rms less than 3.5 mGal, or about twice the rms crossover error (Figure 2). (b) Correlation between the residual gravity anomaly and the topography. In overcompensated or low-density models this is positive while in undercompensated or high-density models this is negative. The change in sign creates a zero contour in this figure; however, no model can achieve a perfect (zero) correlation. (c) The objective function combines the rms and topography correlatig]ns and illustrates that only the model with load density = 2.6 g cm'" and T e = 10 km minimizes both the rms gravity residual and the correlation between the residual gravity and the topography.

Objective Function

Fig. 9. (continued)

we have used the topography of Figure 5, a compromise between our reconnaissance swaths and the map provided to us by N. C. Smoot; if we have underestimated the size of the load, this would explain the high load density implied by Figure 11.

Figure 12 shows a cross section through the best fit model along profile B-B' of Figure 2. From the middle summit to the southeast (to the right) our Sea Beam swath agreed very well with Smoot's map, and the shape of the gravity

SMrrH ET AL.: MAGEI.I. AN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,515

i51o30 '

2½30'

// i i 1

i i.

151o40 '

o ß

ß ß

.. ..

ß ß

, ß

x

/

/ ,. / ..

/ //

\ \\

\

151o50 ' 21o50 '

21o40 '

151ø30 151040' 151 øõO' 1õ2 ø00'

Fig. 10. Map view of residual gravity anomaly at Himu Seamount. There is a small -4-mGal low centered over the northern summit and a broad +8-mGal high centered over the southeast flank.

21o20 '

anomaly is quite sensitive to T e in this region. Visual inspection of model profiles suggests that Te = 15 (+3) lcm, again within the 15% contour of the objective function. Because we are not completely confident that the topography is everywhere correct, we do not feel it is appropriate to make geological interpretations of the residual gravity anomaly at this seamount.

DISCUSSION

Watts et al. [1980] classified Pacific seamounts as having formed "on-ridge" or "off-ridge" using determinations of Te and the empirical observation that T e approximately follows

an isotherm between 300' and 600'C in the cooling plate model [Parsons and Sclater, 1977; Watts, 1978]. They calculated gravity models for T e = 5 km (on-ridge) and 25 km (off-ridge) using a two-dimensional approximation, and they compared gravity profiles along ship tracks to these end-member models. They found that Himu Seamount was better fit by the Te = 5 km model, and Hemler Guyot was better fit by the T e = 25 km model. We have used a full three-dimensional solution with Sea Beam bathymetry and high-resolution gravity; we obtain Te = 10 km at Himu Seamount and 15 km at Hemler Guyot. While our determinations have moved away from the end-members of Watts et al. [1980], the results are in the same sense; we

10,516 SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA

Objective Function

2.8

2.6

2.4 I I I !

0 5 10 15 20 25

Te (km)

Fig. 11. Objective function at Hemlet Guyot. The minimum is not as definitive as at Himu Seamount, probably due to uncertainties in the topography.

find Te at the "on-ridge" seamount is less than at the "off- ridge" seamount, and the 40Ar/39Ar dates also show that the "off-ridge" seamount is younger than the "on-ridge" seamount. We have allowed the load density to be a degree of freedom in our T e determinations, whereas Watts et al. [1980] assumed a fixed value of 2.8 g cm-3. The misfit statistics in Figure 9 suggest that if we had assumed the load density to be 2.8 g cm-3, we also would have found Te = 5 km, confirming their result for Himu Seamount. This would not be the case at Hemler Guyot, where the shape of the guyot strongly violates the symmetry approximation used by Watts et al. [1980].

Recently, it has been recognized that T e values at seamounts in French Polynesia are systematically less than that predicted from the Te-age relation used by Watts et al. [1980]. Calmam and Cazenave [1987] determined Te at 60 volcanic loads worldwide and found that T e values outside French Polynesia follow closely the depth to the 400øC isotherm in the cooling plate model [Parsons and Sclater, 1977], in good agreement with the 450øC isotherm suggested by Watts [1978]. Calmant and Cazenave [1987] also find that young volcanoes in French Polynesia consistently yield Te values less than the depth to this isotherm; they attribute this to a regional thermal anomaly. It is possible that older guyors in the Tuamotu [Cole, 1959; Menard, 1964] and Austral [Menard, 1964; Johnson and Malahoff, 1971] archipelagoes may have biased a few Te determinations in these two areas, but other determinations

from more isolated young volcanoes in these chains also yield unusually low values. In Figure 13 we show the T e values determined in this study in relation to all T e values at seamounts. The distinction between French Polynesian and non-French Polynesian localities is quite clear; the latter yield values which fall between the 300 ø and 500øC cooling

plate isotherms and cluster around a median value of 400øC. This clustering provides an empirical rule of thumb for the relationship between T e and age, and the values from the seamounts of this study and localities in French Polynesia are anomalous with respect to this empirical rule. We infer from Figure 13 that the lithosphere under French Polynesia is unusually warm at shallow levels affecting the elastic portion of the lithosphere and that this was also true in the Cretaceous under the seamounts of this study. It should be noted that Figure 13 includes values from hot spots with large midplate swells, including Hawaii, Cape Verde, and Bermuda, and these data are included in defining the empirical 400øC T e-age rule. This implies that in these localities the excess heat responsible for the midplate swell has not affected the elastic portion of the lithosphere, while in the French Polynesia region the thermal anomaly has penetrated to shallower levels.

Two explanations for the excess heat at shallow levels in French Polynesia have been proposed. Menard and McNutt [1982] suggested that hot spot volcanism produces thermal rejuvenation, in effect resetting the lithosphere temperature to that of a younger plate. McNutt and Fischer [1987] suggested that a thinner lithosphere underlies French Polynesia. Either of these concepts will explain the elevated temperatures in French Polynesia today; the major distinction between these models is only seen as the lithosphere grows old. A rejuvenation event would shift the origin to the right in Figure 13 but would not change the shape of the curves tracking the isotherms, so that after sufficient time the plate temperatures and the seafloor depth would approach their normal asymptotic values. A thinner lithosphere would move the isotherm curves up to shallower levels in Figure 13 without changing the origin, and the asymptotic temperature structure at old age would remain warmer so that a depth anomaly would remain.

The rejuvenation model may be applied to the Te values at our seamounts as follows. We assume that the Te values indicate the depth to the 400øC isotherm in a cooling plate and thereby determine the (rejuvenated) thermal age of the plate when it was loaded by the seamount. At Himu Seamount, T e = 10 km gives a thermal age of 13 m.y. when the seamount was emplaced at 120 Ma, so the lithosphere responded to the seamount load as if the lithosphere had been formed at 133 Ma. Similarly, at Hemler Guyot, Te = 15 km gives a thermal age of 30 m.y. when this seamount was emplaced at 100 Ma, as if the lithosphere had been formed at 130 Ma. These rejuvenated lithosphere ages are quite similar to one another because the age difference between the two seamounts expected from their Te difference [Watts et al., 1980] is very close to their actual age difference determined using 40Ax/39Ar geochronology. The thermal ages for the lithosphere are considerably younger than the 162-165 Ma we infer from the magnetic lineation data (Figure 1). The inferred rejuvenations of 29 and 35 m.y. are schematically shown shifting the 400øC isotherm to the right in Figure 13 (dotted curves).

We have drawn the predicted seafloor subsidence [Parsons and Sclater, 1977] for each of these ages in Figure 14. The curve labeled normal seafloor indicates the subsidence

expected for 163 Ma lithosphere. The curves labeled rejuvenated seafloor track the subsidence expected for lithosphere formed at 133 and 130 Ma but are only drawn from the time of emplacement of each seamount. Above the

SMITH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,517

Gravity - residual

-125 -

225 -

100 -

-25

225 -

L• •oo -

-25 -

oq

3000

6000

density 2.8 T - 15 km

e

observed

Topography

I I I I I I I 0 25 50 75 100 125 150 175

Fig. 12. Profile along B-B' of Figure 2 showing the fit of the gravity model at Hemlet Guyot.

rejuvenated seafloor curves we have added the heights of the seamounts to indicate the subsidence of their summits. This rejuvenation model suggests that the seamounts formed on a depth anomaly of 0.8-1.0 km and that therefore their summits were in shallow water at that time; the model predicts that this depth anomaly has decayed to ~0.15 km today. A model based on thin lithosphere to explain the low T e (e.g., a 75-km plate [McNutt and Fischer, 1987]) would also predict a Cretaceous depth anomaly and place the seamount summits in shallow water, but in this case a large (1.5 km) depth anomaly would persist to the present day. Depth anomalies between 0 and 0.5 km are reported in the region of the Magellan Seamounts [Watts et al., 1985; Cazenave et al., 1986; Cazenave and Dominh, 1987].

We point out that since localities with large midplate swells are included in the empirical Te-age relationship (Figure 13), no thermal rejuvenation can be inferred from Te in these localities. For example, an analysis like that in Figure 14 using data from Hawaii would not predict the existence of the Hawaiian swell, because at Hawaii the thermal age inferred from T e in Figure 13 equals the actual age of the lithosphere at the time of loading. The evidence for a thermal origin for the Hawaiian swell comes chiefly from the decrease in average depth of the swell away from the Hawaiian hot spot [Derrick and Crough, 1978]; the normal Te values require that this excess heat lies below the elastic lithosphere [Von Herzen et al., 1982; McNutt, 1987]. Therefore T e anomalies alone do not uniquely determine

10,518 SMITH ET AL.: MAGEI.!AN SEAMOUNT CRETACEOUS RECORD OF SOPITA

5O

Age at time of loading (m.y.)

0 50 100 150 i

[] .- ß

[][] ..

ß This Study 13 South Pacific

ß Other Localities

I i

0 50 100 150

Fig. 13. Elastic thickness of the lithosphere T e at seamounts and the age of the plate at the time of loading. Data sources are as given by Watts et al. [1982], with additional points from McNutt and Menard, [1978], Watts and Ribe [1984], Fischer et al. [1986], Calmant [1987], McNutt [1988], and Sheehan and McNutt [1989]. Open squares are from localities in French Polynesia, solid circles from localities outside French Polynesia, and solid diamonds indicate the values determined in this study. The shaded area lies in the 300 ø- 500øC region of the cooling plate model of Parsons and Sclater [1977]. The solid circles cluster around the 400øC isotherm (solid curve). Thermal rejuvenation displaces this isotherm to the right (dotted curves).

subsidence histories, and conversely, depth anomalies and subsidence histories do not imply Te anomalies. The actual subsidence history of the Magellan Seamounts could depart from that shown in Figure 14 if there were anomalous heat below the elastic portion of the lithosphere.

However, the flat tops of Hemler Guyot, the hyaloclastites in the dredged materials, and our inferred low-density peaks at Himu Seamount all suggest that the summits of these seamounts were near sea level when they formed, and this would require a Cretaceous depth anomaly of the order of 1 km, independent of the argument based on T e. This amplitude is similar to that of the broad depth anomaly in present-day French Polynesia [McNutt and Fischer, 1987]. Thus these two Magellan Seamounts show evidence for both depth and Te anomalies in the Cretaceous which are similar to those described in French Polynesia today, and the rejuvenated subsidence history suggested in Figure 14 is consistent with the geological and geophysical data.

Menard [1964] analyzed the subsidence of guyors and inferred that a broad region of the western Pacific seafloor had once been much shallower than its present depth. This region, which he called the Darwin Rise, is similar to the present-day depth anomaly in French Polynesia in areal extent [McNutt and Fischer, 1987] and includes the Magellan Seamounts of this study. Menard [1984] reviewed the evidence for formation of the Darwin Rise seamounts at hot

spots in French Polynesia. We can suggest sources in the Cook-Austral islands for the seamounts of this study by "backtracking" using the 40Ar/39Ar ages and the stage poles for Pacific plate motion [Duncan and Clague, 1985; R. (3. Gordon and L. J. Henderson, unpublished manuscript, 1985]. Figure 15 shows these tracks in relation to hotspots active in the Tertiary from Turner and Jarrard [1982] and Duncan

1-

• 3-

q• 4-

r• 5-

Sea Level

• • • Summtts

Himu I . ' ' -- -- -- -- -- -- -- -- --- Seamount

Hemler

Guyot

Normal

Seafloor Rejuvenated

Seafloor

175 150 125 100 75 50 25 0

Depth ]Anomaly

Million years before present

Fig. 14. Subsidence history for the target seamounts and underlying seafloor inferred from the age and T e data of this study. A thermal rejuvenation model to explain the low Te values would imply a Cretaceous paleodepth anomaly; this is independently confirmed by the geology of the seamount summits which were near sea level when they formed. The depth-age relation used is from Parsons and Sclater [1977].

SM1TH ET AL.: MAGELLAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,519

140 ø 30 ø

150 ø 160 ø 170 ø 180 ø 190 ø 200 ø 210 ø 220 ø 230 ø 240 ø 250 ø 260ø 30 ø

20 ø 20 ø

10 ø 10 ø

o o

10 ø 10 ø

20 ø 20 ø

30 ø 30 ø

40 ø 40 ø

50 ø 140 ø 150 ø 160 ø 170 ø 180 ø 190 ø 200 ø 210 ø 220 ø 230 ø 240 ø 250 ø

50 ø 260ø

Fig. 15. Backtracking the seamounts of this study in the Pacific hot spot reference frame. Solid lines are drawn using the stage poles of Duncan and Clague [1985]; dashed lines from poles of R. G. Gordon and L. J. Henderson (unpublished manuscript, 1985). Circles are Tertiary hot spots from Duncan and Clague [1985] and Turner and Jarrard [1982]: HW, Hawaii; SC, Socorro; CA, Caroline; MQ, Marquesas; SM, Smnoa; ME, Mehitia; RA, Rarotonga; RU, Rurutu; PT, Pitcairn; SG, Sala y Gomez; EA, Easter;, MC, Macdonald. Solid lines highlight portions of the East Pacific Rise which are spreading or subsiding asymmetrically [Cochran, 1986]. The track from Himu Seamount passes through the Marshall Islands and ends at Rurutu hot spot; the track from Hemler Guyot passes through the East Mariana Basin and ends at Rarotonga hot spot. Boxed area shown enlarged as Figure 16.

and Clague [1985]. The tracks are drawn from our seamounts at their present ages to the point where the track would be at zero age (the inferred origin of the seamount); the track from Himu Seamount ends near Rurutu and the track from Hemlet

Guyot ends near Rarotonga. There are two sources of error in the backtracking shown

in Figure 15: uncertainties in the age of the seamounts and uncertainties in plate motion. A number of studies have been made of the possible nonfixity of hot spots which might contribute to errors in plate motion. Molnar and Stock [1987] compared the Emperor-Hawaiian track with tracks of hot spots in the Atlantic and Indian oceans and, by assuming that motions between plates were well known, suggested that motions of the order of 10 mm/yr may occur between hot spots under separate plates. If we take this estimate as a maximum for the uncertainty in backtracking, then over the 120 m.y. shown in Figure 15, 1200 km of error might accumulate. Even with this maximum error the Magellan Seamount province can be tied to French Polynesia at the regional scale. Sager and Bleil [1987] assume that the

hot spots within the Pacific form a rigid "frame" and give evidence for motion of this frame with respect to paleomagnetic and paleontologic latitude indicators; the very success of their analysis indicates that motions among hotspots within the Pacific basin must be small. Watts et al. [1988] have shown that stage poles for the Emperor- Hawaiian chain [Epp, 1978] give a good fit to the Louisville Ridge; the misfit between Emperor-Hawaiian and Louisville poles amounts to 2 ø of displacement in 70 m.y. Using this as an estimate of intra-Pacific hotspot drift gives 300 km total error in Figure 15, one fourth of that implied by the interplate study of Molnar and Stock [1987].

A larger source of error is the uncertainty in seamount age. We believe that the 40Ar/39Ar geochronology presented here accurately dates the rocks dredged from these seamounts, but it is conceivable that the bulk of the seamounts formed

at some time other than the rocks we have dated. The

petrology of these rocks is consistent with the later stages of volcanism described in the Hawaiian Islands [Clague and Dalrymple, 1987], where Recent lavas are found at Kauai and

10,520 SMrrH E'r AL.: MAG• SEAMOUNT CRETACEOUS RECORD OF SOPITA

185 ø 5 ø

o

10 ø

15 ø

20 ø

25 ø

:30 ø

35 ø 185 ø

190 o 195 ø 200 ø

•3o o

205 ø 210 ø 215 ø 220 ø

ratak

225 ø

MQ

himu ' •s• •

230 ø

hemler PT•}

o ø erikub •

MC ß

0

.

190 ø 195 ø 200 ø 205 ø 210 ø 215 ø 220 ø 225 ø 230 ø

235 ø 5 ø

10 ø

15 ø

20 ø

25 ø

30 ø

35 ø 235 ø

Fig. 16. Qualitative error estimate for backtracking. Hot spot locations are as in Figure 15. Shown are inferred source locations for Hemler Guyot and Himu Seamount and also for Erikub and Ratak guyots of the Marshall Islands [Davis et al., 1989]. Six possible source locations are shown for each seamount (triangles); ellipses summarize how errors in age and plate motion affect inferred source location. The "errors" are created by using the two different plate motion models of Figure 15 and adding _+ 5 m.y. to the ages determined from dredged rocks. Hemler Guyot yields an inferred source position near present-day Rarotonga hot spot and is isotopically similar to Rarotonga. Himu Seamount and Erikub and Ratak guyots are inferred to have come from a source near Rumtu and are isotopically similar to one another and to Rumtu.

Niihau, 5 m.y. "downstream" of the present location of the Hawaiian hot spot [Clague and Dalrymple, 1988]. We therefore will take 5 m.y. as an estimate of the age error in backtracking. In Figure 16 we have computed six possible source positions for each seamount: three using the plate motion model of Duncan and Clague [1985] and three using that of R. G. Gordon and L. J. Henderson (unpublished manuscript, 1985). Each set of three corresponds to three ages: the actual rock age we determined for each seamount and 5 m.y. more and less than this age. The six source locations for each seamount are summarized with ellipses. If the locations were drawn at random from a bivariate normal

distribution, then the ellipses would enclose the source regions with 95% confidence; however, since we do not have a statistical model for the backtracking error, the ellipses are not quantitative error ellipses in the strict sense. If the bulk of the seamounts is older than the rocks we have dated, as

the Hawaiian growth model suggests, then the source of the seamounts would be in the eastern halves of the ellipses.

The ellipses in Figure 16 suggest that Himu Seamount formed near the present-day location of Rurutu and Hemlet Guyot near the present-day position of Rarotonga. Hot spots have been proposed at these locations to explain the age progressions in the Cook-Austral Islands [Turner and

Jarrard, 1982]. The isotope data from Himu Seamount plot in the Rurutu field, and the isotopic character of rocks from Hemlet Guyot is similar to that of Rarotonga (Figure 7). This geochemical coincidence lends support to these backtracked source positions. Additional substantiation may be provided by recent data from the Marshall Islands archipelago. The Himu Seamount track passes through the Marshall and Gilbert islands (Figure 15), and Staudigel et al. [1987, 1989 submitted], and Davis et al. [1989] have reported isotopic signatures from Erikub and Ratak guyots very similar to those of Himu Seamount and Rurutu Island (Figure 7). We have backtracked Erikub and Ratak Guyots in Figure 16 in the same manner as Hemlet Guyot and Himu Seamount; Ratak Guyot yields a source location identical to that of Himu Seamount and Rurutu Island. Erikub Guyot does not seem to have come from any known Tertiary hotspot, but if it were -7 m.y. younger than the age given by Davis et al. [1989], then it too could have formed at a source near present-day Rurutu Island. We feel that these data from the Marshall archipelago support the backtracking of our two Magellan seamounts and indicate that high-resolution geochronology now permits tracing the movement of individual seamounts. We do not mean to imply, however, that all seamounts in the Magellan and Marshall areas were

SM1THET AL.: MAGEI.IAN SEAMOUNT CRETACEOUS RECORD OF SOPITA 10,521

formed by Cook-Austral hot spots; our analysis does not preclude other volcanic events at other times contributing to a complex history in these regions [Schlanger and Moberly, 1985; Schlanger et al., 1987].

There are a variety of unusual features which characterize the South Pacific today (Figure 15). These include radiogenic isotope anomalies in young volcanoes [Dupr• and All•gre, 1983; Hart, 1984; Palacz and Saunders, 1986], slow seismic velocities at deep [Dziewonski and Woodhouse, 1987] and shallow [Nishimura and Forsyth, 1985] levels, asymmetric spreading and subsidence at the East Pacific Rise [Cochran, 1986], small-scale convection [Haxby and Weissel, 1986], a widespread depth anomaly [McNutt and Fischer, 1987], low elastic strength of the lithosphere [Calmam and Cazenave, 1987], and a large number of active volcanoes. Menard [1984] referred to these volcanoes as the "Polynesian Plume Province," but there is some evidence that the thermal anomaly may begin "upstream" of the French Polynesian volcanoes, at or near the East Pacific Rise [Cochran, 1986; Haxby and Weissel, 1986]. McNutt and Fischer [1987] called the depth anomaly the "South Pacific Superswell," but their region does not include the Marquesas Islands which exhibit the isotope and elastic thickness anomalies characteristic of this region. The Hawaii, Bermuda, and Cape Verde hot spots produce midplate swells, but T e is not unusual in those localities; it is the elastic thickness anomaly, rather than the swell, which distinguishes this area of the South Pacific. We use the name South Pacific Isotopic and Thermal Anomaly (SOPITA) to refer to this association of anomalies in depth, elastic thickness, and radiogenic isotopes. Staudigel et al. [1987] used the term SOPIA in their discussion of the isotope anomalies.

Some of the characteristic features of the SOPITA today may not be preserved in the geologic record of an ocean basin. Recognition of ancient asymmetric spreading requires access to portions of the conjugate plates which are now subducted. Ancient thermal anomalies would presumably decay so that small-scale convection and seismic slowness anomalies would disappear. The only way to infer ancient thermal anomalies is through the subsidence history of guyots and the elastic thickness of the lithosphere which records the thermal state of the plate at the time of seamount loading. Isotope anomalies in the rocks would of course be preserved. Thus we have probably found in the Magellan Seamounts all the evidence of an ancient SOPITA which can

be observed.

There is one additional feature which the present SOPITA and the Early Cretaceous Magellan Seamounts have in common and that is a large westerly component of plate motion with respect to the hot spots [Duncan and Clague, 1985]. Note in Figures 15 and 16 that during times of predominantly westerly plate motion (such as the last 43 m.y.) the lithosphere passes over several melt sources in turn. Lithosphere that was over the Macdonald melt source in the Oligocene was near the Rurutu source in the Miocene is now near Rarotonga and will later pass Samoa. Similarly, the lithosphere in our seamount study area passed over a source of HIMU magma at 120 Ma and 20 m.y. later passed a source of EM-11 material. Menard [1984] reviewed the idea that the Darwin Rise was created by the lithosphere passing over several hotspots. Note in Figure 15 that if plate motion "tracks" were drawn from the SOPITA hot spots the

tracks would converge on the Darwin Rise region in the Early Cretaceous, because prior to 100 Ma the plate motion also had a large westerly component. Between 100 and 43 Ma the motion was more northerly, separating the effects of the melt sources.

In the SOPITA today it is not clear whether the many hot spots are a cause or an effect of the thermal anomalies. If the conflagration of melt sources which occurs with westerly plate motion is a cause of the SOPITA thermal anomalies, then we should find these anomalies reduced or absent in

seamounts formed between 100 and 43 Ma. Therefore

seamounts formed during this period in chains such as the Marshall Islands may provide a measure of the role of plate motion in the SOPITA phenomena. A discussion of the geologic h/story of the Marshall Islands is beyond the scope of this paper and can be found elsewhere [cf. Schlanger et al., 1987]. We point out here that some of the SOPITA isotopic characteristics have been found at Erikub and Ratak guyors [Staudigel et al., 1987; Davis et al., 1989], and the subsidence history of the region suggests a Cretaceous palcodepth anomaly [Menard, 1964; Schlanger and Premoli- Silva, 1981; Schlanger and Moberly, 1985; Schlanger et ai., 1987]. On the other hand, Watts et al. [1980] found that Te in this region was better fit by a large value (25 kin) than a small value (5 km). We have seen in this paper that a seamount which Watts et at. [1980] thought better fit by the large value (Hemlet Guyot) in fact has an intermediate value which is anomalously low in the manner of $OPITA seamounts. We speculate that some seamounts formed during periods of northerly plate motion may be found to exhibit SOPITA characteristics. We hope that future studies using accurate gravity, geochronology, and isotope data from seamounts along these hot spot tracks will look for the Te, palcodepth, and isotope anomalies of the SOPITA.

CONCLUSIONS

We have analyzed geophysical and geochemical data from two Cretaceous seamounts in the western Pacific which were

formed at melt sources in what is now French Polynesia. The flexural compensation of these seamounts indicates an anomalously low elastic thickness consistent with Te data from French Polynesia. The T e values found by three- dimensional analysis of BGM-3 gravity and Sea Beam data, 10 (ñ 2) and 15 (ñ 3) km, are closer together than the 5 and 25 "end-members" assigned to these seamounts by the two- dimensional reconnaissance study of Watts et al. [1980] but are of the same sense; the "on-ridge" seamount (Himu Seamount) gives a lower T e than the "off-ridge" seamount (Hemler Guyot). The age difference between the seamounts determined by 40Ar/39Ar geochronology is consistent with that inferred from their T e difference and an empirical 400øC isotherm rule relating Te to (rejuvenated) thermal age. Thermal rejuvenation of T e with respect to this empirical rule is a phenomenon described only at these two Magellan Seamounts and those of French Polynesia, and we suggest that this indicates that the lithosphere in these regions carries excess heat at shallower levels than lithosphere under "classical" hot spots such as Hawaii. These Cretaceous seamounts also display isotopic characteristics very similar to young basalts from French Polynesia, indicating that the mantle isotope end-members HIMU and EM-II have been closely associated spatially for at least 100 Ma and

10,522 SMITH ET AL.: MAGla7.! AN SEAMOUNT CRETA•OUS RECORD OF SOPITA

suggesting that the origin of each is linked to the same geological process. The isotope, T e, and depth anomalies. which characterize the South Pacific Isotopic and Thermal Anomaly today are all displayed in the seamounts of our study, and we conclude that the SOPITA was active in the Early Cretaceous when it produced the Darwin Rise. If the SOPITA phenomena are caused by plate motion inducing hot spot conflagration, then seamounts formed during the northerly plate motion of 100-43 Ma may not display these phenomena. Combined geophysical and geochemical investigations in the Marshall-Gilbert archipelagoes and the East Mariana Basin may show whether the SOPITA phenomena have been produced continuously since the Cretaceous. In either case, the SOPITA phenomena are long lived.

Acknowledgements. The data analyzed in this paper were collected during 5 days of ship time funded by the Office of Naval Research under contract N00014-84-C-1032, Scope WSP. Additional gravity data analysis was supported in part by the Office of Naval Research under contract N00014-87-K-0204, Scope I. Isotope analyses were supported by National Science Foundation grant OCE-87-11798. We thank R. S. Detrick, M. Diament, J. Francheteau, D. G. Martinson, W. B. F. Ryan, and P. Wessel for helpful discussions. The manuscript benefitted from the comments of the JGR Associate Editor and also from reviews by T. M. Brocher, A. Cazenave, M. A. Lanphere, J. L. Rubenstone, and S. O. Schlanger. Lamont-Doherty Geological Observatory contribution 4494.

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W. H. F. Smith and A. B. Watts, Lamont-Doherty Geological Observatory, Columbia University, Palisades, NY, 10964.

H. Staudigel, Scripps Institution of Oceanography, University of California, San Diego, La Jolla, CA, 92093.

(Received November 29, 1988; revised May 23, 1989;

accepted May 25, 1989.)