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Keywords:

  • seismic imaging;
  • intraoceanic arc;
  • Izu-Bonin;
  • crustal structure;
  • subduction zone;
  • rifting

Abstract

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

Crustal evolution processes in intraoceanic arcs, including crustal accretion and rifting, have been long discussed. To examine crustal evolution in the Izu-Bonin intraoceanic island arc, we conducted an active-source wide-angle seismic study along a north-south profile (500 km long) within a possible paleoarc in the rear arc (i.e., the Nishi-shichito ridge) about 150 km west of the present-day volcanic front. In this study, the seismic velocity and reflectivity images are obtained using the wide-angle seismic data. For the seismic velocity imaging, we applied refraction tomography in which 93,535 picks were used. The overall root-mean-square (rms) misfit calculated from the initial model of the refraction tomography was 483.1 ms, and those calculated from the final model were reduced to 66.7 ms. The resultant seismic image shows marked variations of crustal thickness along the seismic profile: thin crust (10–15 km thick) in the northern part, three discrete thick crustal segments (20–25 km thick) in the central part, and a moderately thick crust (∼15 km thick) in the southern part. These variations are mainly attributed to thickness variations of the middle crust having seismic velocity of 6.0–6.8 km/s. This variation of crustal thickness does not correlate with seafloor topography, which is characterized by post-Miocene across-arc seamount chains. It does correlate well with crustal variations observed along the present-day volcanic front of the Izu-Bonin arc. These findings suggest that the magmatic activity that created the across-arc seamount chains had little effect on the rear-arc crust and that the main part of the rear-arc crust was created before the rear arc separated from the volcanic front. By correlating the structural variations along the rear arc (i.e., the variation of the average seismic velocity as well as the thickness of the middle crust) and those along the present-day volcanic front, we found that the direction of rifting to separate the rear arc (paleoarc) from the present-day volcanic front was north-northeast.

1. Introduction

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

The Izu-Bonin-Mariana (IBM) arc is an intraoceanic island arc at the eastern edge of the Philippine Sea plate where the Pacific plate is subducting beneath the arc. An intraoceanic arc such as the IBM arc provides an excellent opportunity to examine the process of evolution of new crust because an intraoceanic island arc is less affected by preexisting continental crust than one at the edge of a continent. Previous petrological studies [e.g., Taylor, 1967, 1977] have proposed that post-Archean growth of andesitic continental crust was mainly accomplished by accretion of island arc crust onto continental crust (the andesitic model). Understanding the processes of generation of new island arc crust is, therefore, fundamental to the examination of the processes by which continental crust develops on the present-day Earth.

Seismic velocity models of arc crust have provided strong constraints on the petrological models for growth of island arc crust. Suyehiro et al. [1996] presented the first seismic image to cross an entire active intraoceanic arc and observed a layer with a P wave velocity (Vp) in the range 6.0–6.4 km/s in the middle of the crust beneath the volcanic front. Because these velocities are within the range of velocities in the upper continental crust [Christensen and Mooney, 1995], Suyehiro et al. [1996] and Takahashi et al. [1998] concluded that felsic to intermediate plutonic rocks were created in the middle of the crust even though large basalt volcanoes were observed in the Izu arc. Another characteristic of the northern Izu arc is a high-velocity layer (Vp = 7.1–7.3 km/s) in the lower crust. The high-velocity layer represents about 35% of the total crust beneath the volcanic front [Suyehiro et al., 1996; Takahashi et al., 1998]. Similar structures have been observed at other intraoceanic arcs; for example, the Tonga ridge [Crawford et al., 2003] and the Mariana arc [Lange, 1992; Takahashi et al., 2007, 2008; Calvert et al., 2008].

On the other hand, different seismic structures have been observed for the Aleutian arc. A seismic model by Holbrook et al. [1999] shows no evidence of a layer with Vp of 6.0–6.4 km/s but has layers with Vp of 4.3–5.4 and 6.5–6.8 km/s in the upper and middle crust, respectively. They suggested that these two layers may represent a remnant of original oceanic crust on which the island arc was built. Holbrook et al. [1999] also identified a thick high-velocity layer (10–20 km thick) at the lower crustal level. Higher pressure in the lower crust due to the thicker arc crust (∼30 km thick) may explain the higher Vp, but it should be noted that the velocity beneath the volcanic front (Vp = 6.9–7.3 km/s) was slightly slower than that of the Izu arc. To transform the present-day Aleutian crust into a mature continental crust, Holbrook et al. [1999] proposed that two additional processes would be necessary, namely, intracrustal melting to form an intermediate-composition middle crust and delamination of the mafic lower crust.

Recent seismic studies along the volcanic front of the Izu-Bonin arc [Kodaira et al., 2007a, 2007b] provide new seismological constraints on the growth of island arc crust. Kodaira et al. [2007b] showed that (1) continental crust with Vp of 6.0–6.8 km/s was generated predominantly beneath basaltic volcanic centers, and (2) the bulk composition of the crust does not change from the thin Bonin arc crust in the south to the thick Izu arc crust in the north, and it is more mafic (basaltic) than typical continental crust. Those observations suggest that for arc crust to evolve into continental crust, there must be a process to return mafic to ultramafic lower crustal components to the mantle.

Rifting and back-arc spreading processes are also important when considering crustal evolution in an intraoceanic arc. The IBM arc has been under strong arc normal extension [Stern et al., 2003], which has provided the driving force to create a back-arc basin as well as a remnant arc at the edge of the back-arc basin. In the IBM arc system, only the older crustal components are, in general, preserved in the west. On the other hand, a complete history of crustal evolution is preserved around the present-day volcanic front [e.g., Okino et al., 1999; Stern et al., 2003]. Comparison of the crustal structure of the volcanic front and that of the remnant arc, therefore, provides information about crustal growth processes after the remnant arc was separated.

In this study, we used wide-angle seismic data (Figure 1) to determine the structure of the crust and uppermost mantle along the Nishi-shichito ridge, ∼150 km west of the present-day volcanic front (hereafter we refer to the area between the volcanic front and the Shikoku basin around the Nishi-shichito ridge as the rear arc). We then develop the crustal evolution process for the Izu-Bonin arc by comparing the structure of the rear arc with that of the volcanic front, which was imaged previously by Kodaira et al. [2007b] in a study of the present-day volcanic front.

image

Figure 1. Maps of study area showing the regional setting and location of the wide-angle seismic profile (red line). The wide-angle seismic profile is located along the Nishi-Shichito Ridge ∼150 km west of the present-day volcanic front. The black line shows the location of a wide-angle seismic profile of a previous study [Kodaira et al., 2007b]. Black circles connected by dotted lines show pairs of crustal segments in the rear arc and volcanic front. A–E indicates the crustal segments of the rear arc that are discussed in detail in the text. The blue line shows the MCS profile shown in Figure 2. Hcj, Hachijo-jima; Ags, Aoga-shima; Sms, South Sumisu; Tsm, Torishima; SFG-TL, Sofugan Tectonic line; Tnp, Tempo seamount; Omc, Omachi seamount.

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2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

The IBM intraoceanic arc system extends over 2800 km from Sagami Bay in the north to Guam in the south. Its rifting and back-arc spreading history has been well documented on the basis of petrological studies and geomagnetic studies [e.g., Stern and Bloomer, 1992; Bloomer et al., 1995; Kobayashi et al., 1995; Okino et al., 1999; Hall, 2002; Stern et al., 2003]. Here, we briefly summarize those studies of relevance to rifting and spreading processes.

The age of basement rock sampled from the fore-arc region suggests that initial subsidence of the lithosphere along the IBM arc began at about 45–50 Ma [Bloomer et al., 1995; Cosca et al., 1998; Ishizuka et al., 2006a]. Subsidence of the lithosphere evolved into subduction at about 43 Ma (Eocene) when the motion of the Pacific plate changed abruptly from a northward to a westward direction [Richards and Lithgow-Bertelloni, 1996]. Single-arc volcanism close to its present position continued until rifting started along the Parece Vela Basin at 30 Ma (Oligocene) in the southernmost part of the IBM. The Shikoku Basin started spreading at the northernmost end of the IBM arc at 25 Ma. The two rifting systems met at 20 Ma and spreading stopped at about 15 Ma. Consequently, the Kyushu-Palau Ridge (KPR) was formed at the western edge of both basins. Back-arc rifting along the Mariana Trough started at 10 Ma and seafloor spreading began at 3–4 Ma [Bibee et al., 1980; Yamazaki et al., 2003]. The above tectonic history along the Mariana arc suggests that the conjugate pair of the KPR is preserved as the West Mariana Ridge. On the other hand, a conjugate pair of the KPR in the Izu-Bonin arc is unclear. There has been disagreement whether it is at the western margin of the Shikoku Basin [e.g., Kobayashi and Nakada, 1978] or east of the present-day volcanic front [e.g., Shiki, 1985; Chamot-Rooke et al., 1987].

Detailed magnetic studies, however, have demonstrated that the counterpart of the KPR is at the rear arc. Okino et al. [1994] and Yamazaki and Yuasa [1998] found north-south aligned dipole anomalies or long-wavelength magnetic lineations (138.5°–139°E to the north of 27°N), which show similar anomalous patterns to those observed along the KPR. The magnetic anomalies were interpreted to represent induced magnetization associated with middle- to lower-crustal plutonic bodies that were formed before or during the opening of the Shikoku Basin [Yamazaki and Yuasa, 1998].

In addition to the long-wavelength magnetic anomalies observed at the rear arc, Yamazaki and Yuasa [1998] found another north-south trending long-wavelength magnetic lineation immediately east of the volcanic front. They concluded that the two magnetic anomalies provide evidence of back-arc rifting of early Miocene age. Taylor [1992], however, concluded that the entire Izu arc was stretched during the Oligocene, creating an extensional regime in the crust between the fore arc and rear arc before the opening of the Shikoku Basin. However, it must be emphasized that although there are some inconsistencies concerning the timing of rifting between the fore arc and the rear arc (i.e., before or after the Shikoku Basin opened), the above three studies [Taylor, 1992; Okino et al., 1994; Yamazaki and Yuasa, 1998] agreed that the conjugate part of the KPR in the Izu-Bonin arc is beneath the rear arc.

Systematic temporal and spatial variations of volcanic activity after spreading of the Shikoku Basin stopped have been reported for the area around the rear arc [Ishizuka et al., 2002]. The seafloor topography and composition of volcanic rocks allows this area to be divided into three zones from west to east: a zone of across-arc seamount chains, a zone of back-arc knolls, and an active rift zone [e.g., Ishizuka et al., 2002]. The currently active rift zone is 10–15 km west of the volcanic front. There is Holocene basaltic lava (<0.1 Ma) at the central axis of the rift, and older lava (∼1.4 Ma) has been sampled at the rift wall [Hochstaedter et al., 1990a, 1990b; Urabe and Kusakabe, 1990; Ishizuka et al., 2003, 2006b]. The back-arc knoll zone consists of north-south trending ridges and small volcanoes, most of which are 500–1000 m in diameter [Honza and Tamaki, 1985]. The volcanism in this zone had been active from 2.8 to 1 Ma [Ishizuka et al., 2003]. The across-arc seamount chains are a distinctive feature of the Izu part of the IBM arc; they consist of en echelon, northeast-southwest trending ridges. Volcanism along the across-arc seamount chains began at 17 Ma and continued until 3 Ma [Ishizuka et al., 1998, 2003]. Ishizuka et al. [2003] clearly demonstrated that volcanism in the three zones migrated from west to east after 17 Ma. According to the tectonic process described above, we concluded that volcanism after the Miocene may have been superimposed onto a crust formed before the opening of the Shikoku Basin in the Oligocene.

3. Data Acquisition

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

A 500-km-long profile of wide-angle seismic data along the rear arc from 27.5°N to 32°N was acquired in June 2006 by R/V Kairei of the Japan Agency for Marine-Earth Science and Technology (JAMSTEC) (Figure 1). Details of the cruise are provided by Kodaira et al. [2006]. A total of 100 ocean-bottom seismographs (OBSs) were deployed at approximately 5-km intervals along the profile. The OBSs were positioned on the seafloor using a supershort baseline acoustic positioning system, and their positions were further calibrated by using the traveltime of direct water waves. All OBSs were recovered; however, five of them (OBS09, 25, 30, 45, and 52) malfunctioned and did not record usable signals. The OBSs used (4.5-Hz, three-component gimbal-mounted geophones and hydrophones, continuous 16-bit digital recording with 100 Hz) were originally designed by Kanazawa and Shiobara [1994] and Shinohara et al. [1993]. An air gun array on R/V Kairei was used at a shooting interval of 200 m, which corresponds to a time interval of ∼90 s. The air gun array comprised eight Bolt long-life air guns with a total volume of 12,000 cubic inches (197 L). The standard air pressure was 2000 psi (140 MPa). A differential GPS was used for positioning the shots. The accuracy of the location of each shot was about 0.4 m.

Multichannel seismic reflection (MCS) data acquired along the wide-angle profile and adjacent profiles acquired by an earlier cruise of R/V Kairei [Tsuru et al., 2005; Kaiho et al., 2006; No et al., 2007] were used to construct the shallow section down to acoustic basement (see Figure 2), which was used as an initial model for the seismic refraction tomography of this study. The seismic reflection data were recorded with a 204-channel hydrophone cable (∼5100 m). The air gun system as described above was used to acquire the MCS data, but the shooting interval was 50 m, corresponding to a time interval of ∼20 s.

image

Figure 2. Time-migrated MCS section along the profile shown in Figure 1 (blue lines), showing the (left) EW and (right) NS sections. The north-south section coincides with part of the wide-angle seismic profile.

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4. Wide-Angle Seismic Data

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

Wide-angle seismic data from every fifth OBS (20 km interval) are shown in Figure 3, except for OBS25, OBS30, and OBS45, which did not record usable signals (OBS26, OBS31, and OBS44 are shown instead). Arrival times were strongly affected by the variation of water depth along the profile. The application of static corrections for the water layer greatly improved phase identifications. We applied simple preprocessing to the wide-angle seismic data, including a 5–15 Hz minimum phase band-pass filter to increase the signal-to-noise ratio, and a predictive deconvolution filter to reduce reverberations after strong first and later arrivals. Automatic gain control (3-s window) was used to enhance identification of first arrivals.

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Figure 3. Examples of wide-angle seismic data. (a) OBS05, (b) OBS10, (c) OBS15, (d) OBS20, (e) OBS26, (f) OBS31, (g) OBS35, (h) OBS40, (i) OBS44, (j) OBS50, (k) OBS55, (l) OBS60, (m) OBS65, (n) OBS70, (o) OBS75, (p) OBS80, (q) OBS85, (r) OBS90, (s) OBS95 and (t) OBS100. The reduction velocity used was 8 km/s. The horizontal axis indicates the offset distance from the OBS displayed. A 5–15 Hz band-pass filter and 3-s automatic gain control were applied. (bottom) Picked refraction traveltimes (red lines), calculated refraction traveltimes from the final model (yellow line), and picked reflection traveltimes (blue line) are plotted.

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Figure 3. (continued)

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Figure 3. (continued)

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Figure 3. (continued)

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Figure 3. (continued)

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Data quality was generally good. First arrivals were identified for offsets greater than 100 km for most of the data (e.g., Figures 3a, 3b, 3e, 3g, 3h, 3i, 3k, 3l, 3n, and 3o), but, for some OBSs, first arrivals were identified only up to offsets of several tens of kilometers (e.g., Figure 3c, 3j, and 3s). We interpreted refraction first arrivals to represent several crustal and mantle phases. Although phase identification of the refraction first arrivals is unnecessary for the modeling procedure we used (i.e., refraction tomography), in order to describe the observed data, we classified the refraction arrivals into two distinct phases: those with an apparent velocity of less than 8 km/s and those with an apparent velocity of about 8 km/s. For convenience, we refer hereafter to the former as the crustal refraction phase and to the latter as the mantle refraction phase. We were able to infer a general pattern of structural variation directly from the observable offsets of the first arrivals of the crustal refraction phase and from the intercept times of the mantle refraction phase.

The first arrivals of the crustal refraction phase as are evident over only a short range of offsets (∼20–50 km of shot-receiver distances, Table 1) for the OBSs in the northern part of the profile (OBS5 to OBS31). In the middle part of the profile (100–400 km from the northern end of the profile), those ranges are larger (more than 100 km of shot-receiver distances) than in other parts of the profile, except for OBS40 and OBS65, which show a shorter observable range of crustal refraction phases (less than 50 km of shot-receiver distances). The OBSs at the southern end of the profile (OBSs 85–100) show moderate observable ranges of the crustal refraction phase (80–140 km of shot-receiver distances). Although, as stated above, we divided the refracted phases into two distinct phases on the basis of a critical apparent velocity of about 8 km/s, none of the data shows the mantle refraction phase to have apparent velocities markedly higher than 8.0 km/s. The velocities of all of the mantle phase refractions were slightly slower than 8.0 km/s. A mantle phase of similar velocity character was observed by Kodaira et al. [2007a] in the data from the volcanic front.

Table 1. Maximum Offsets of Crustal Refraction Arrivals
Ocean-Bottom SeismographOffset (km)
0520–40
1020–40
1540
2040
2640
3140–50
35100
40120–140
44120
5060?
55100
60100
6550
70100
7580
8080?
8550
9090
9560
10060?

The examples of the refraction data (Figure 3) show several wide-angle reflection phases of complicated character; i.e., it was difficult to correlate those phases for each OBS. Although the modeling procedure employed here for the reflection phases does not require correlation of each observed reflection, we interpreted them as belonging to two groups: a reflection phase approaching a branch point of crustal and mantle refractions and a reflection phase observed at further offsets than the branch point. The former group is interpreted as reflections from the base of the crust. Some of the latter group approach the mantle refraction phase and are interpreted as reflections from a reflector in the mantle. The reflections interpreted to be from the base of the crust are strong and clearly observed at some OBSs, for example, 80–100 km offsets at OBS35 and OBS40. But no clear reflections from the base of the crust are observed at some of other OBSs, for example, on the southern side of both OBS75 and OBS100. Although the lack of those reflections may be a function of data quality, it may also suggest that there is only a small velocity contrast between crust and mantle for some parts of the profile. The reflection phases approaching mantle refraction phases are observed at far offsets in many OBSs: for example, 160–180 km offset (i.e., shot-receiver distance) at OBS05, at offsets greater than 150 km at OBS15, greater than 200 km at OBS20, greater than 160 km at OBS70, and offsets of 160–180 km at OBS90. These observations indicate that there are reflectors of only local extent within the mantle.

5. Modeling Procedure

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

To quantitatively compare seismic images of the volcanic front and the rear arc, we used the same modeling procedures that we used in previous studies [Kodaira et al., 2007a, 2007b]. The procedure consists of two steps, seismic velocity imaging and seismic reflectivity imaging. For the seismic velocity imaging, we used refraction tomography. Recent wide-angle seismic studies have demonstrated that a tomographic approach is more objective and robust than the conventional layer-stripping approach [e.g., Zelt et al., 2003]. The software we used for refraction tomography was GeoCT-II, which is based on the algorithm of Zhang et al. [1998]. We constructed a velocity model with 0.5 × 0.5 km grid for forward calculation and inversion of first arrivals to obtain a seismic velocity image. For traveltime and raypath calculations, the software uses a wavefront method [Zhang and Toksöz, 1998] employing graph theory that allows only straight rays within a cell of constant slowness [Zhang et al., 1998]. In the inversion step, the Tikhonov regularization is applied and an objective function containing three terms associated with the misfit of average slownesses, the misfit of apparent slownesses and the roughness of model slowness is minimized [Zhang and Toksöz, 1998]. To define the roughness of a model in the objective function, the second-order derivative of model slowness is used [Zhang and Toksöz, 1998]. There is a tradeoff parameter between model misfit (the first two terms of the objective function) and model roughness (the third term of the objective function). Generally, with the tradeoff parameter increasing, the misfit of traveltime decreases and model roughness increases [Zhang et al., 1998]. A few trials of the inversion are needed to select an optimal tradeoff parameter, which allow the inversion to converge to a predetermined data misfit residual [Zhang and Toksöz, 1998]. In our case, we selected the tradeoff parameter to reduce the overall root-mean-square (rms) misfit to less than uncertainties of picked traveltimes and to avoid geologically meaningless structural roughness. Details of the traveltime calculations and inversion methodology are provided by Zhang et al. [1998].

The seismic velocity image we obtained is expressed as a continuous velocity field that does not generate reflection phases, but strong wide-angle reflection phases were observed in the OBS data, as shown in section 4. In order to include this reflectivity information, we used a diffraction-stack migration approach using picked reflection traveltimes [Ito et al., 2005; Fujie et al., 2006]. The principle of this method is the same as that for diffraction stack of wide-angle prestack depth migration [McMechan and Fuis, 1987; Schleicher et al., 1993; Simon et al., 1996], but only diffraction contours corresponding to picked reflection traveltimes are projected onto the velocity structure obtained by refraction tomography. As demonstrated by Fujie et al. [2006], this procedure is a more effective way to obtain lithospheric-scale reflectors from wide-angle reflection data than is a conventional diffraction-stack method. In the conventional method, all waveforms are projected along diffraction contours onto a velocity structure, which often provides a poor image because several phases, such as refraction phases, converted phases, and multiple reflection phases, overlap with the target reflection phases.

A final velocity image was obtained in three steps as follows. First, a laterally uniform structure (excluding the part of the section shallower than the acoustic basement obtained by the depth conversion of the MCS data using the results of velocity analysis of the MCS data) (Figure 4a) was constructed as an initial model for the first tomographic inversion (Figure 4b). In order to construct the initial model, we referred to the velocity structure of the southern Izu arc obtained by our previous study [Kodaira et al., 2007b].

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Figure 4. (a) Initial model for the first tomographic inversion. A laterally uniform velocity field was used except structures shallower than acoustic basement which was obtained from MCS data (Figure 2). (b) Results of the first inversion. (c) Initial model for the second tomographic inversion. The model was obtained by smoothing and simplifying the results of the first inversion. (d) Results of the second inversion. A reflectivity image obtained from a diffraction-stack migration using picked traveltimes is superimposed. (e) Initial model for the final tomographic inversion. Abrupt changes of the velocity gradients were introduced at the 7.0 and 7.6 km/s isovelocity contours on the basis of strong reflectors observed on the reflectivity image. Overall root mean squared misfits calculated from each model are indicated above the figures.

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Two sets of arrival time data were used in this step. Near offset arrivals (less than 30 km offset) were used to constrain the shallow section, and all observed arrival time data were used to construct a model for the deeper section. In the second step, the final model of the previous step (Figure 4b) was simplified and smoothed, and then used as the initial model (Figure 4c) for a second tomographic inversion. The diffraction-stack migration approach using picked reflection traveltimes was performed by using the final velocity model for this inversion (Figure 4d). Although strong continuous reflectors are not clearly evident, there are locally continuous reflectors around isovelocity contours at ∼7.0 and ∼7.6 km/s. We therefore included abrupt velocity increases along the ∼7.0 and ∼7.6 km/s isocontours to construct an initial model (Figure 4e) for the final inversion. In the final step, the diffraction-stack migration approach was applied for all picked reflection phases by using the final velocity model (Figures 5 and 6).

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Figure 5. (a) Seafloor topography along the profile. The names of the seamounts close to the profile are labeled. (b) Final seismic velocity image. The shaded area indicates the poorly resolved area identified by the checkerboard test (Figure 7). Annotations A–E show thicker parts of the crustal segments. (c) Geological interpretation of the final seismic velocity image with seismic reflectors from shallower than the 7.8 km/s isovelocity contour overlaid.

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Figure 6. Seismic reflectivity image deeper than the 7.8 km/s isovelocity contour is superimposed on the geological interpretation of the final seismic velocity image.

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6. Seismic Velocity and Reflectivity Images

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

6.1. Traveltime Fitting and Resolution

We used 93,535 picked first arrival data for the refraction tomography. The overall rms misfit calculated from the initial model of the first inversion step (Figure 4a) was 483.1 ms. The overall rms misfit calculated from the final model (Figure 5b) became 66.7 ms after 10 iterations. Other overall rms misfits calculated from models of each inversion step are indicated in Figure 4. Traveltime residuals of each OBS calculated from the final model are shown in Figure 7a, and calculated and picked traveltimes are superimposed on the observed data in Figure 3. The traveltime residuals of each OBS were less than ±0.05 s (Figure 7a), which is comparable to the uncertainties of the arrival-time data observed at most OBSs. Traveltime misfits calculated from the final model for each shot at each OBS (Figure 7b) show that the misfits for the shots even at larger offsets are lower than 0.05 s for most shots.

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Figure 7. (a) Mean traveltime residuals for each OBS calculated from the final velocity model. Vertical bars show the standard deviation. (b) Traveltime residuals for each shot at each OBS. The vertical axis shows the OBS number (see Figure 1). The color scale indicates the traveltime residuals for each shot. (c) Result of checkerboard test. In the original checkerboard pattern, velocity anomalies of up to 5% were defined by a 10 km (horizontal) × 5 km (vertical) grid at depths shallower than 10 km and by a 20 km × 10 km grid below 10 km. The area above the dotted line is well resolved.

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Checkerboard tests are commonly used to assess model resolution in tomographic studies [e.g., Hearn and Ni, 1994: Zelt and Barton, 1998]. In this study we used the checkerboard test described by Zelt and Barton [1998]. Synthetic data from a known model consisting of the initial model of the final inversion with an anomaly pattern of positive and negative regions were inverted using the same method as that used to obtain the final model. The velocity anomaly pattern is a sin(x)sin(y) function with a peak anomaly value of ±5% of the background velocity. The source and receiver geometry of the field survey was used, and the data were inverted using the same initial model. On the basis of ray coverage, we expected higher resolution in the shallow part of the profile. We thus introduced different anomaly sizes for different depths. For depths less than 10 km we used a 10 (horizontal) × 5 (vertical) km grid. For depths greater than 10 km we used 20 (horizontal) × 10 (vertical) km grid.

The checkerboard test (Figure 7c) shows good recovery of the pattern down to 20 km along the entire profile. At depths of 20–30 km, although the amplitude of the pattern weakens, the positive-negative patterns are still recognizable, except in the northern part of the profile. These results substantiate the structural variation described in the next section, that is, large-scale structural differences in the northern, central, and southern parts of the profile, as well as the three-segment structure of wavelength 80–100 km in the central part of the profile. Our modeling procedure did not allow quantitative estimation of uncertainty, but we could deduce the uncertainty at particular depths from the recovered checkerboard peak values. For example, in the central part of the profile, the 5% peak anomaly was recovered by the checkerboard test as a ∼3% peak anomaly, corresponding to an uncertainty of ∼0.1 km/s for a velocity of 7 km/s, which was the velocity observed at depths of 15–20 km. Thus, the velocity uncertainty at 15–20 km depth was ∼0.1 km/s. We expect that uncertainties would be even smaller for depths shallower than 20 km, where the recovered checkerboard patterns were clearer.

6.2. Seismic Images

The final seismic velocity and reflectivity images are shown in Figures 5 and 6. For convenience, we divided the model into three layers, primarily on the basis of velocities, as we did in our previous studies of the volcanic front [Kodaira et al., 2007a, 2007b]. The three layers were those with Vp < 5.8 km/s, Vp = 6.0–6.8 km/s, and Vp = 6.8–7.6 km/s. Hereafter, we refer to these layers as the upper, middle, and lower crust, respectively. We further subdivided the lower crust into two layers separated by the 7.2 km/s isovelocity contour. The partially observed clear continuous reflectors were also taken into account in defining the layer boundaries. Although a recent petrological study that used seismic images from the IBM arc [Tatsumi et al., 2008] interpreted a layer with Vp = 7.4–7.6 km/s as part of the mantle consisting of crustal component (i.e., mafic cumulates in the mantle), in this study we followed the layer nomenclature of our previous study [Kodaira et al., 2007a, 2007b].

In the northern part of the profile (0–120 km), the model shows a thin (10–15 km thick) crust of relatively uniform thickness (Figure 5). We recognized a slight broadening of the isovelocity contours and downward convex reflectors, both of which represent crustal thickening, to a maximum depth of 15–17 km deep beneath the Kanbun (at 45 km) and Jokyo (at 90 km) seamounts. Although the model suggests crustal thickening beneath the seamounts, the thickness change is remarkably smaller (less than ∼5 km) than that observed in the central part of the profile (described in the following paragraph). The MCS section (Figure 2) shows bright reflectors at about 2 s below the seafloor, corresponding ∼3 km below the seafloor assuming an average velocity of 3 km/s above the reflector, between the seamounts. These reflectors consist of several subparallel reflectors of lower frequency and different character to those from the top of the Shikoku Basin crust. We believe, therefore, that the reflectors at about 2 s below the seafloor between the seamounts do not represent the top of the Shikoku Basin but may correspond to the top of the crust of an old arc system.

In the central part of the profile (120–370 km), the model shows strong lateral variations of crustal thickness. Here, there are three segments, each of which is about 80–100 km wide. The thickest parts are at about 150, 250, and 330 km on the profile, where each segment shows a crustal thickness of 20–25 km. These variations are consistent with the structural character inferred from the observed variations of the offset range of crustal refraction arrivals (Table 1). The seismic image (Figure 5) shows that the structural variation is attributed mainly to the middle crust (Vp = 6.0–6.8 km/s), which thickens to 7–10 km toward the center of each crustal segment but becomes remarkably thin (less than 2 km thick) between segments (at 200 and 290 km). The upper part of the lower crust (Vp = 6.8–7.2 km/s) shows a similar pattern; that is, 5–7 km thickness in the center of each segment and as thin as the middle crust between segments. There is not a strong lateral thickness variation in the lower part of the lower crust (Vp = 7.2–7.6 km/s); it is 5–7 km thick though the entire central part of the profile. It is important to note that the crustal structure does not correlate with seafloor topography, which is characterized by the across-arc seamount chains (Figure 5a). The modeled crustal structure shows three distinct segments in the central part of the profile where neither the size nor the spacing of the seamounts corresponds to those three segments. Although the reflectors that may correspond to the base of the crust are recognized beneath the southern two segments around the 7.6 km/s isovelocity contour, many of the reflectors are discontinuous and thus are not indicative of laterally continuous layer boundaries along the entire profile (Figure 5).

In the southern part of the profile (370–500 km), the model shows a structure that is intermediate between those of the northern and central parts of the profile. The crust thickens (∼17 km thick) near the Kita-Kyowa and Bunka seamounts. As was apparent in the center of the profile, crustal thickening at the Bunka seamount (440 km) is attributed mainly to thickening of the middle crust (Vp = 6.0–6.8 km/s). The reflectors are not clearly imaged in this part of the profile; however, the reflectors imaged at 17–18 km depth beneath the Bunka seamount may correspond to the base of the crust.

As previously described, we observed far offset reflection phases that we interpreted to be reflections from within the mantle. We plotted separately (Figure 6) the mantle reflectors (those deeper than the 7.8 km/s isovelocity contour) to better distinguish them from the crustal reflectors (Figures 5c). Velocities in the mantle at depths greater than 35 km are not well controlled; however, by assuming a mantle velocity of ∼8 km/s we identified clear reflectors at 35–40 km depth in the center of the profile in addition to the reflectors at 25 km depth (Figure 6). Reflectors in the mantle at around 40 km depth have also been observed along the IBM volcanic front [Kodaira et al., 2007a; Takahashi et al., 2008; T. Sato et al., Amplitude modeling of the seismic reflectors in the crust-mantle transition layer beneath the volcanic front along the northern Izu-Bonin island arc, submitted to Geochemistry, Geophysics, Geosystems, 2008].

7. Discussion

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

7.1. Possible Paleoarc Structure Deduced From Seismic Imagery and Magnetic Anomalies

The volcanic front of the Izu-Bonin arc is believed to have been close to its present position since the Oligocene [e.g., Stern et al., 2003]. This means that the crust around the current volcanic front has preserved the entire process of crustal formation since the Oligocene. On the other hand, a remnant arc (or paleoarc) that has been separated from the volcanic arc by rifting preserves only crustal formation processes for the period before rifting occurred. Crustal evolution before and after rifting can, therefore, be deduced by comparing the structure of the current volcanic front with that of a paleoarc.

It is generally accepted, on the basis of magnetic data[e.g., Kobayashi et al., 1995; Okino et al., 1999], that the KPR is a remnant Oligocene arc that has been separated from the IBM arc by the opening of the Shikoku and Parece Vela basins, which began at 25–30 Ma. On the other hand, the Oligocene arc crust in the Izu-Bonin arc is still being debated. Most previous studies have suggested that the Oligocene crust in the Izu-Bonin arc lies western edge of the arc beneath the rear arc [e.g., Taylor, 1992; Okino et al., 1994; Yamazaki and Yuasa, 1998], but a few studies have concluded that the Oligocene crust is east of the present-day volcanic front [e.g., Shiki, 1985; Chamot-Rooke et al., 1987]. Yamazaki and Yuasa [1998] suggested that there is Oligocene crust in the rear arc that was separated from the volcanic front by rifting in the early Miocene (i.e., after opening of the Shikoku Basin) on the basis of north-south trending long-wavelength magnetic anomaly lows. We hereafter examine whether or not the magnetic anomalies proposed as evidence of paleoarc crust represent crustal-scale structures.

We plotted our seismic profile on part of the magnetic anomaly map of Yamazaki and Yuasa [1998] (Figure 8). The original map was a total-force magnetic anomaly map of the northern part of the Philippine Sea plate that was compiled by Yamazaki et al. [1991], Ishihara [1994], and Ishihara and Kishimoto [1996]. The magnetic anomaly map is a reduced-to-pole map, continued upward to 12 km, and was constructed as follows [Yamazaki and Yuasa, 1998]. First, a total-force magnetic anomaly map was reduced to the magnetic pole; that is, the total field anomaly was transformed into an anomaly that would be measured at the north magnetic pole [Blakely, 1995]. The reduced-to-pole transformation projects a dipole-type total force anomaly in middle latitudes as a positive anomaly. Second, the reduced-to-pole anomaly map was continued upward to 12 km. This transformation is effectively a spatial high-cut filter that reduces the effect of short-wavelength anomalies caused by shallow sources. The above transformations tend to emphasize long-wavelength anomalies, which may be attributed to deep sources at middle crustal level [Yamazaki and Yuasa, 1998].

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Figure 8. The wide-angle seismic profile is plotted on part of a magnetic anomaly map from Yamazaki and Yuasa [1998]. Annotations a–e indicate peaks of the magnetic anomalies immediately east of the seismic profile. STL, Sofugan Tectonic line; SKR, Shin-Kurose ridge. White dots show locations of the north-south lineation of magnetic anomalies interpreted by Yamazaki and Yuasa [1998].

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The westernmost of three north-south alignments of long-wavelength magnetic anomalies observed by Yamazaki and Yuasa [1998] corresponds to the KPR. Of the two other north-south aligned magnetic anomalies, one corresponds to the rear arc and the other lies close to, but slightly east of, the current volcanic front (Figure 8). The seismic profile of our study lies 20–50 km west of the magnetic anomaly along the rear arc, which shows five clear magnetic highs (marked with a–f in Figure 8). We found a remarkably good correlation between the seismic velocity image and the arrangement of magnetic highs. The three northern strong magnetic highs (marked with c–e in Figure 8) are immediately east of the three thick crustal segments (marked with C–E in Figure 5), where the middle crust (Vp = 6.0–6.8 km/s) thickens down to 15 km depth. The two smaller magnetic highs (marked with a and b in Figure 8) correspond to the slightly thicker crust at 400 and 450 km on the profile (marked with A and B in Figure 5). It is also notable that the crust is thin at the northern part of the profile (0–120 km) where the profile is situated 50 km apart from a broader and weaker positive magnetic anomaly (marked with f in Figure 8).

From these observations, we concluded that the north-south alignment of magnetic anomalies can be attributed to the crustal-scale structural variation which is mainly due to the variation of the middle crust between 5 and 15 km depth. This is consistent with the interpretation of Yamazaki and Yuasa [1998] who suggested the presence of magnetized bodies within the middle crustal level. We also concluded that similar structural variations may exist that correspond to the north-south trending magnetic anomalies observed at the KPR. In other words, our observation suggests the alignment of the middle crust having felsic-to-intermediate component along the volcanic front, the rear arc, and the KPR. We believe that existence of the paleoarc crust separated from the volcanic front by rifting process is the most likely interpretation of the alignment of the felsic-to-intermediate component crust along the rear arc as well as the KPR.

We believe that the above considerations provide new evidence that demonstrates the presence of paleoarc crust along the rear arc. In our previous study, the middle crust and the upper and lower parts of the lower crust were interpreted as felsic-intermediate plutonic rocks, mafic plutonic rocks, and a mixed zone of mafic-ultramafic rocks partly composed of olivine cumulates, respectively, based on the results of sonic wave measurements for plutonic rocks sampled at Tanzawa, Japan [Kitamura et al., 2003]. If the rear arc consists of crust that has been separated from the volcanic front, we would expect the individual layers within the rear arc to have similar compositions to those of the volcanic front. It is underlined that it should be careful to discuss structural variations of across arc direction (i.e., from the rear arc to the volcanic front) on the basis of the magnetic anomaly data because the curie depth map around Japan region [Okubo et al., 1989] shows that curie depth shallows eastward from the rear arc to the volcanic front to less than 7 km depth at the west of the volcanic front.

7.2. Structure Along the Rear Arc and Volcanic Front

As discussed above, the seismic image we obtained strongly supports the presence of a paleoarc crust (possibly Oligocene crust) beneath the rear arc. Comparing the structure of the present-day volcanic front with that of the rear arc, therefore, will provide insight into the processes of crustal growth in the Izu-Bonin arc. Our previous study [Kodaira et al., 2007b] showed along-arc variations of crust to demonstrate important structural characteristics that constrain the growth of island arc crust. Those are the variation pattern of the average seismic velocity of crust which reflects bulk chemical composition of crust [e.g., Smithson et al., 1981; Kelemen and Holbrook, 1995; Shillington et al., 2004] correlates with the distribution of the basalt volcanoes, and the felsic-to-intermediate component middle crust having seismic velocity of 6.0–6.8 km/s is thickened toward centers of each basalt volcanoes.

To see if there are similar variations along the rear arc, we calculated the average seismic velocity of the crust and the thickness of the middle crust with Vp of 6.0–6.8 km/s (Figure 9). To calculate the average seismic velocities, we used the same criteria as those used in our previous study [Kodaira et al., 2007b]. That is, the average velocities were estimated by using the seismic velocity between the top of the middle crust and the bottom of the lower crust. Crustal material with velocity slower than 6.0 km/s (the upper crust) was excluded because upper crustal velocities are strongly affected by several parameters other than crustal composition, such as variable fracture distribution and porosity [e.g., Carlson and Gangi, 1985; Kelemen and Holbrook, 1995].

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Figure 9. Curves showing lateral variations of average seismic velocity of crust and thickness of the middle crust. (top) Seafloor topography. Annotations A–E show thicker parts of the crustal segments as shown by the seismic image (Figure 5).

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The variations of average seismic velocity and thickness of middle crust along the rear arc show similar patterns to those along the present-day volcanic front, except at the northern end of the rear-arc profile where thin crust was imaged. The variations of both average velocity and thickness of the middle crust clearly show a wavelength of 50–80 km between 100 and 500 km along our profile (Figure 9). Similar patterns are also observed along the present-day volcanic front in the Izu arc [Kodaira et al., 2007b]. Although the maximum thicknesses of the middle crust at segments A to E (4–10 km thick, Figures 9) is one half to two thirds of those beneath the basaltic volcanoes of the present-day volcanic front, the average velocities beneath the thickest parts of segments A to E are almost identical to those beneath the basaltic volcanoes (∼6.8 km/s). This suggests that the volume of crust for each segment of the rear arc is smaller than those at the basalt volcanoes, but the bulk compositions of the crust are almost identical. We once again emphasize that these variations of the rear arc do not correlate with seafloor topography, which is characterized by across-arc seamount chains created by magmatic activities after the Miocene [e.g., Ishizuka et al., 2002]. This suggests that the magmatic activity that created the across-arc seamount chains had little effect on the rear-arc crust, and that the main body of crust at the rear arc was formed before it separated from the volcanic front.

Comparison of the velocity-depth profiles of the rear arc with those of the volcanic front also supports the above interpretation. In our previous study [Kodaira et al., 2007b], we showed that vertical extension of the velocity-depth profile beneath the Suiyo seamount, which is a basaltic volcano on thin crust of the Bonin arc, produced a velocity-depth profile similar to that beneath Aoga-shima in the Izu arc, where the crust is thick. We plotted the velocity-depth profile beneath segment D of the rear arc with those at basaltic volcanoes of the volcanic front (Figure 10). The velocity-depth profile of segment D lies between those of the Suiyo seamount and Aoga-shima, and the 150% vertical extension of the segment D profile below 6 km/s is almost identical to that of Aoga-shima. This suggests that the crust beneath the basaltic volcanoes of the present-day arc has evolved by continuous thickening while maintaining the volume ratios of each crustal component at least after the rear arc separated from the volcanic front.

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Figure 10. One-dimensional profile of segment D of the rear arc plotted with velocity-depth profiles at Aoga-shima and the Suiyo seamount in the volcanic front [Kodaira et al., 2007b]. The dashed black line shows the 150% vertically extended profile at segment D below the 6 km/s isovelocity contour, and the dashed blue and red lines show 150% and 250% vertically extended profiles beneath the middle crust at Aoga-shima and Suiyo seamount, respectively [Kodaira et al., 2007b]. The velocity-depth profiles of typical continental structures compiled by Christensen and Mooney [1995] (C&M) and Rudnick and Fountain [1995] (R&F) are superimposed.

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Although, the average velocities beneath the thickest part of segments A to E are almost identical to those beneath the basaltic volcanoes, the average velocity between the segments in the rear arc shows slightly higher velocity (>7.1 km/s) (Figure 9) than those between the basaltic volcanoes (6.9–7.1, but mostly less than 7.0 km/s) [Kodaira et al., 2007a, Figure 11]. This difference is not significant, but seems to be mainly attributed by two factors concerning the lower crust; i.e., higher volume ratios and higher velocities of the lower crust between the segments in rear arc. Since lower temperatures in the rear arc away from the volcanic front are expected, the observed velocity difference may reflect a temperature difference. However, we could not discuss this quantitatively, due to lack of heat flow data around our profiles.

If we assume that the locations of the sources of magma that built the arc crust have changed little since the rear-arc crust separated from the volcanic front, identifying conjugate pairs of crustal segments in the rear arc and the present-day volcanic front will gain an understanding of the rifting or extensional process and its direction. To pursue this, we examined the correlation between the structural variations along the rear arc (i.e., the spatial pattern of the average velocity of the crust as well as the thickness of the middle crust) and those along the volcanic front (Figure 11). It should be noted that the structural variations in the northern part of the rear-arc profile (dotted lines in Figure 11) were ignored when we examined correlation, because we believe that the main part of the paleoarc may be situated more eastward of the profile, since our profile was taken obliquely to the magnetic lineation. In Figure 11, we considered three cases of correlations of the structural variations; that is, we plotted three versions of the variation curves with segment E of the rear arc aligned in turn with Hachijo-jima, Aoga-shima, and Sumisu-jima. Of these, the alignment with Hachijo-jima shows the best correlation (Figure 11a). Here, segments A, C, D, and E of the rear arc correspond to Tori-shima, Sumisu-jima, Aoga-shima, and Hachijo-jima, and the peak between Minami-Sumisu caldera and Tori-shima corresponds to segment B. It is noteworthy that this comparison not only shows matching peak-and-trough patterns, but the shapes and slopes of the curves are also similar. The other two comparisons (Figures 11b and 11c) lack a correlation of the peak-and-trough patterns south of Sumisu-jima. Even in areas where the peak-and-trough patterns correlate reasonably in Figures 11b and 11c, the shapes of the curves do not match as well as those of Figure 11a. Each of the pairs of structural segments we matched in Figure 11a is connected by lines in Figure 1. The north-northeast trend of these lines is parallel to the Sofugan Tectonic line and to the trend of topographic highs on the seafloor between the rear arc and volcanic front (Figures 1 and 12). Therefore, we concluded that the rear-arc crust has a paleoarc structure that has been separated from the volcanic front by north-northeasterly extensional rifting. This idea is supported by a study which deduced ages of seamounts at both end of the Sofugan Tectonic Line [Ikari, 1991]. Ikari [1991] concluded that the age of the Tempo seamount (Figure 1) situated at the southern end of the Sofugan Tectonic Line was comparable with the age of the Omachi seamount (Figure 1) situated at the northern end the Sofugan Tectonic Line (Figure 1). Since the patterns of structural and velocity variation of the paleoarc and the present-day volcanic front are identical, we deduced that the locations of the basaltic volcanoes of the volcanic front changed little before or after the rear arc became separated.

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Figure 11. Comparison of structural variations along the rear arc with those of the volcanic front. The location of the profile along the volcanic front is shown in Figure 1. To examine spatial correlations, the variations of the rear arc have been horizontally shifted so that segment E corresponds to (a) Hachijo-jima, (b) Aoga-shima, and (c) Sumisu-jima. The red and black curves show the thickness of the middle crust and the average seismic velocity of the crust in the volcanic front, respectively [Kodaira et al., 2007b]. The orange and blue curves show the thickness of the middle crust and the average seismic velocity of the rear arc, respectively. Hcj, Hachijo-jima; Ags, Aoga-shima; Sms, South Sumisu; Trs, Torishima.

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Figure 12. Three-dimensional block diagram with seismic images of both the rear arc and the volcanic front [Kodaira et al., 2007b] showing the direction of rifting as suggested in this study (gray arrows). SFG-TL, Sofugan Tectonic line.

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The reflectors observed in the uppermost mantle at 35–40 km deep have been recently reported at several areas at a volcanic front and a remnant arc in the IBM arc [e.g., Kodaira et al., 2007a; Takahashi et al., 2008]. Although discussion about an origin of the deep reflectors is beyond a scope of this study, we prefer to an interpretation described by Takahashi et al. [2008]; that is, the possible origin of those reflectors might be formed the transformation of the mafic/dense crustal material brought by repeated crustal growth.

8. Conclusions

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

In this study, we used wide-angle seismic data to model the structure of the crust and uppermost mantle of the rear arc that lies on the western edge of the Izu-Bonin intraoceanic arc. Previous geophysical and geological studies have proposed that the rear-arc crust may represent a paleoarc crust (presumably of Oligocene age). We investigated the origin of the rear-arc crust and considered the evolution of the crust of the intraoceanic arc by comparing the structure of the rear arc with that of the present-day volcanic front.

Seismic image along the rear arc showed strong lateral variations that are mainly attributed to thickness variations of the middle crust having seismic velocity of 6.0–6.8 km/s. In the northern part of the profile, the model showed thin crust (10–15 km thick) that thickened slightly under seamounts. In the central part of the profile, we identified three distinct segments of crust, each 80–100 km wide and with maximum thicknesses of 20–25 km. The crust in the southern part of the profile showed a structural character intermediate between those of the northern and central parts of the profile. The crust thickens (to ∼15 km) close to seamounts.

We found a good correlation between seismically defined structural variations and a long-wavelength magnetic anomaly pattern that has been interpreted to represent a paleoarc that forms a conjugate pair to the KPR. The strong peaks of the magnetic anomalies at the rear arc are in areas where we found thicker middle crust. Those observations add new evidence to demonstrate the presence of a counterpart crust to that of the KPR (i.e., a presumably Oligocene paleoarc) along the rear arc. The variations of average seismic velocity reflecting bulk composition of crust and the thickness of the middle crust along the rear arc show a similar pattern to those along the present-day volcanic front, except at the northern end of the rear arc profile where the crust is thin, but the structural variations within the rear arc do not correlate with seafloor topography, which is characterized by across-arc seamount chains created after the Miocene. These findings suggest that the magmatic activity that created the rear-arc seamounts had little effect on the rear-arc crust and that the main part of the rear-arc crust was created before the rear arc separated from the present-day volcanic front. By correlating the structural variations along the rear arc (i.e., the variation of the average seismic velocity as well as the thickness of the middle crust) and those along the present-day volcanic front, we found evidence that the extensional regime of rifting that formed the rear arc was likely oriented north-northeast.

Acknowledgments

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References

This study was funded by the Institute for Research on Earth Evolution, Japan Agency for Marine-Earth Science and Technology, and partially supported by Grant-in-Aid for Creative Scientific Research (19GS0211) and Grant-in-Aid for Scientific Research (B) (20340122). We gratefully acknowledge the captain, crew, and technical staff on board the R/V Kairei for their help with acquisition of the seismic data.

References

  1. Top of page
  2. Abstract
  3. 1. Introduction
  4. 2. Tectonic Setting Relevant to Rifting and Spreading in the IBM Arc
  5. 3. Data Acquisition
  6. 4. Wide-Angle Seismic Data
  7. 5. Modeling Procedure
  8. 6. Seismic Velocity and Reflectivity Images
  9. 7. Discussion
  10. 8. Conclusions
  11. Acknowledgments
  12. References