crustal structure of the mid-atlantic ridge at 23°n from seismic refraction studies

$: Crustal Structure of the Mid-Atlantic Ridge at 23°N From Seismic Refraction Studies$
JOURNAL OF GEOPHYSICAL RESEARCH, VOL. 91, NO. B3, PAGES 3739-3762, MARCH 10, 1986

Crustal Structure of the Mid-Atlantic Ridge at 23øN From Seismic Refraction Studies

G. M. PURDY

Woods Hole Oceanographic Institution, Woods Hole, Massachusetts

ROBERT S. DETRICK

Graduate School of Oceanography, University of Rhode Island, Narragansett

An explosive seismic refraction experiment using four ocean bottom hydrophone receivers was carried out along a 120-km-long section of the Mid-Atlantic Ridge (MAR) median valley at latitude 23øN immediately south of its intersection with the Kane Fracture Zone. The data are interpreted in terms of laterally homogeneous horizontally layered structures using travel time inversion and synthetic seismogram modeling techniques. These solutions are combined using two-dimensional ray tracing to produce a single model defining the major along-axis structural changes occurring beneath this section of the median valley. With the exception of slightly low layer 3 velocities, the crust beneath several tens of kilometers of the MAR median valley has all the characteristics of simple, mature ocean crust including a well-defined Moho transition zone, ~ 8 km/s upper mantle velocities, and a total crustal thickness of 6-7 km. No evidence is found for the presence of a steady state axial magma chamber in the crust or upper mantle, and the excellent propagation characteristics and orderly amplitude versus range relationships imply the existence of a considerable degree of lateral homogeneity beneath the median valley over distances of several tens of kilometers along-axis. We infer from these observations that the basic seismic characteristics of the ocean crust are "frozen-in" in a time period less than that separating the major volcanic injection events (10,000-50,000 years) by hydrothermal circulation penetrating to the base of the crust. Along the 100-km-long ridge segment studied here, two major along-axis structural changes occur. The first is a 10- to 15-km-wide (along-axis) zone of lowered velocities in the lower crust centered beneath a major along-axis topographic high. This is interpreted to be the remnants of the most recent phase of injection that has temporarily left behind it a region of elevated temperatures and pervasive cracking and thus reduced velocities. The second occurs abruptly at latitude 23ø15'N and is accompanied by both a major change in rift valley and crestal mountain morphology and the apparent noncoincidence of the median valley and central magnetic anomaly. North of this boundary the ~30-km-long ridge segment that abuts the Kane Fracture Zone has a crustal thickness of 4-5 km and no distinctive layering to produce the characteristic amplitude patterns that typify mature oceanic crust. Interpretation of the cause of this structural change is uncertain but may be related to local tectonic events, perhaps a recent 10- to 20-km ridge jump to the east.

INTRODUCTION

The most prominent tectonic feature of the Atlantic Ocean is the Mid-Atlantic Ridge (MAR). It extends for over 12,000 km from Iceland to the Bouvet Triple Junction and is a slow spreading ridge (1-3 cm/yr half rate) that has produced in the central Atlantic Ocean the oldest continuous record of sea-

floor spreading in existence in the world's oceans. The processes of accretion of oceanic lithosphere at a slow spreading ridge, and particularly the MAR, have been the subject of many research efforts [e.g., Aumento et al., 1970; Heirtzler and van Andel, 1977]. The determination of the seismic structure of the crust and upper mantle beneath the MAR is particularly important because of its potential to constrain models of the principal thermal, mechanical, and volcanic processes that control lithosphere accretion. Review of existing seismic refraction data •Purdy and Ewing, 1985] shows that our knowledge of this structure is poor and is limited primarily to three small areas (at latitudes 37øN, 45øN, and 60øN).

The first extensive seismic experiment dedicated to the determination of MAR structure was that of Keen and Tramon-

tini •1970] at 45øN. This experi•nent produced one of the largest and most rigorously interpreted data sets ever collected over the MAR but was located 30-40 km west of the median

Copyright 1986 by the American Geophysical Union.

Paper number 5B5508. 0148-0227/86/005 B- 5508505.00

valley and found a mean crustal structure that was essentially normal with an average crustal thickness of 5 km and an upper mantle velocity of 7.9 km/s. The first work utilizing fixed ocean bottom instruments and closely spaced air gun shots to study the structure beneath the median valley itself was that of Whitmarsh [1973, 1975] at 37øN (FAMOUS area [Heirtzler and van Andel, 1977]). These experiments determined widespread anomalously low layer 3 velocities of 6.2 km/s, and a 2- to 3-km-wide axial zone with low layer 2 velocities of 3.2 km/s (presumed to be the zone of intrusion) outside of which the upper mantle velocity was poorly constrained to be 8.1 _+ 0.4 km/s. In a series of papers, Fowler [1976, 1978] and Fowler and Keen [1979] analyzed refraction data from both 37øN and 45øN using reflectivity waveform modeling techniques [Fuchs and Muller, 1971]. Two well- constrained conclusions of this work were that a substantial

magma chamber cannot exist at shallow depths beneath the MAR median valley and that the formation of a "normal" 6- to 7-km-thick crustal section with a 7.2 km/s basal layer and an 8.1 km/s upper mantle occurs within 10 km of the axis. Bunch and Kennett [1980] report a particularly thorough interpretation of a reversed line located on the crest of the Reyk- janes Ridge at latitude 60øN, where a small (0.2 km/s) velocity inversion is found within layer 3 and the upper mantle velocity is as low as 7.1 km/s.

The first observations of a crustal low-velocity zone beneath a spreading ridge were made by Orcutt et al. [1976] and Ro- sendahl et al. [1976] on the East Pacific Rise. No comparable

3739

3740 PURDY AND DETRICK: MID-ATLANTIC RIDGE SEISMIC STRUCTURE

observations that would support the existence of a magma chamber beneath the MAR have been made to date.

Most of the seismic data collected over the MAR has been

located over young crust on the ridge flanks and provides little insight into the structure of the accretion zone itselfi Because of the very rapid cooling of newly formed crust by hydrothermal circulation, the structures associated with crustal accretion processes (as opposed to those simply of young but fully formed crustal sections) will be restricted to a narrow zone only a few kilometers wide. Only some of the data in the four studies by Whirmarsh [1975], Fowler [1976, 1978], and Bunch and Kennett [1980] were suitably located within a sufficiently well-determined tectonic framework to provide some coverage of the accretion zone itselfi From this brief summary it is clear that very little seismic data exist upon which to base a meaningful study of the accretion zone of the MAR. A second source of information concerning accretion processes at slow spreading ridges is the models constructed to satisfy primarily geological observations of ophiolite sequences, the petrology of drilled and dredged rock samples, and what little is known of magma chamber dynamics and processes [e.g., Bryan and Moore, 1977; Pallister and Hopson, 1981; Sleep and Rosendahl, 1979; Stakes et al., 1984].

The section of the MAR that is the subject of this paper is located south of the Kane Fracture Zone (KFZ), a right- lateral transform that offsets the MAR by 150 km and 10 m.y. near latitude 24øN [Fox et al., 1969; Purdy et al., 1978, 1979]. This fracture zone and the adjacent ridge segments have been the subject of intensive investigation [e.g., Detrick and Purdy, 1980; Louden and Forsyth, 1982; Karson and Dick, 1983; Cor- mier et al., 1984; Toomey et al., 1985]. The presence of anomalously thin (2-3 km) crust beneath the Kane Fracture Zone was first reported by Detrick and Purdy [1980] and later confirmed as a characteristic of a 300-km length of the fracture zone by Cormier et al. [1984]. Cormier et al. also described constraints on the complex structure of the southern ridge- fracture zone intersection and presented travel time observations compatible with a thinning toward the fracture zone of the crust beneath the southern median valley. Toomey et al. [1985] interpret well-determined foci and source mechanisms of microearthquakes beneath the median valley 100 km south of the KFZ as evidence that to a depth of at least 7-8 km the lithosphere has cooled to temperatures within the brittle field of behavior and that this particular ridge segment is undergo- ing active extension at this time.

In this paper we interpret new data pertaining to the seismic structure of the crust and upper mantle beneath the center of the median valley along a 100-km-long segment of the MAR centered at approximately 22ø50'N (Figures 1 and 2). Unlike many of the early MAR seismic experiments, this work is focused on the central median valley over the presumed location of the accretion zone. Following a description of the data, the interpretation will be divided into three sections. The first two are concerned respectively with the travel time inversion and synthetic seismogram modeling of the data recorded south and north of a major structural discontinuity that exists at latitude 23ø15'N. These solutions are in terms of horizon-

tally layered laterally homogeneous structures. Features of the record sections and travel time data sets, which cannot be interpreted under these assumptions, are ignored. The third section uses these laterally homogeneous models as the foundation for a more complex two-dimensional structure that attempts to account for all the primary travel time and amplitude characteristics of the data set. The result is a well- constrained model for the seismic structure beneath a 100-km-

long sections of the MAR median valley.

EXPERIMENT LOCATION AND DESCRIPTION

The four Woods Hole Oceanographic Institution ocean bottom hydrophones (OBH) [Koe[sch and Purdy, 1979• used to receive and record the seismic refraction data that are the

subject of this paper were located within the MAR median valley as shown in Figures 1-3 (see Table 1). The explosive shooting lines, located as nearly as possible to the center of the median valley, extend 130 km south of the MAR-KFZ intersection. The median valley is well-developed, and no major transform offsets are identifiable from the relatively good coverage of conventional wide-beam echo-sounding data. Multi-narrow-beam echo-sounding data collected since this experiment was carried out also reveal no morphologica! evidence for the presence of transform offsets along this section of the MAR [Derrick et al., 1984]. There are, however, a number of notable along-strike changes in the ridge morphology. The depth of the median valley r•aches a prominent minimum 70 km south of the KFZ at latitude 22ø55'N (Figure 3). South of this latitude the rift valley walls are noticeably asymmetric, the steepest walls being on the eastern side (Figure 2). North of the median valley depth minimum, near 23ø10'N, the western rift mountains are noticeably subdued, the median valley appears to broaden, and the central magnetic anomaly (or strictly its causative body) is shifted by 10-20 km west of the center of the rift valley [Purdy et al., 1978' Derrick et al., 1984]. Based on a regional study of bathymetry and magnetics, Schouten et al. [1985b] predict the presence of small or zero offset transform zones [Schouten and White, 1980], as shown by the shaded bands in Figure 3' these predicted locations are near deeps in the median valley, and the southernmost zone coincides with a 6- to 7-km right- lateral "bend" in the eastern wall. Detrick et al. [1984-1, however, based on Sea Beam data, report that topographic trends within the inner rift valley are undisturbed across these features.

The ocean bottom hydrophone instruments are numbered 1, 2, 3, and 6' OBH 1 is located on the previously mentioned median valley high at 22ø55'N (Figure 2) at a depth of 3217 corrected meters and, as can be seen in Figure 3, is flanked by a 1.2 ø average dip down northward to OBH 2 and a 1.8 ø average dip down to the south. OBH 2 is located approximately 30 km to the north of OBH 1 at a depth of 3958 corrected meters, just a few kilometers south of the northern limit of the (apparent) central axial high. OBH 3 is located 62 km north of OBH 1 at what was considered to be the north-

ernmost limit of the identifiable median valley at a depth of 4200 corrected meters. Ten kilometers farther north, OBH 6 is positioned at a depth of 4702 corrected meters on the edge of the nodal deep which marks the intersection of the KFZ and the MAR. Results pertaining to the structure of the intersection of the KFZ and the median valley obtained as part of this experiment are described by Cormier et al. [1984].-

Three shooting tracks were carried out over these four receiving instruments, once firing 14.5 kg (32 lb.) T. NT charges every 1 km, a second time firing 112.6 kg (256 lb.) TNT charges every 3 km, and finally a 1000 in. 3 air gun line with 100-m shot spacing. The mean detonation depth for the 112.6 kg charges was 66 ___ 4 m and for the 14.5-kg charges 51 ___ 1.5 m.

Data Reduction

Firing time corrections (typically, 0.3 s for the 112.6 kg, 0.11 s for the 14.5 kg, and, of course, zero for the air gun) were computed in the standard manner and applied to the seismograms, as were corrections for receiver clock drift (typically less than 10 ms). Corrections for seafloor topography were

PURDY AND DETRICK' MID-ATLANTIC RIDGE SEISMIC STRUCTURE 3741

25

24

47W ß .

46 45 44 43

45 44 43

::::::::::::::::::::::::::::::::

ß

::':::'i!

42

25

24

23

22

42

Fig. 1. Simplified bathymetric map of the Kane Fracture Zone and the Mid-Atlantic Ridge. The MAR median valley is indicated by the diagonal lines. Depths greater than 4000 m are densely stippled and the 3000- to 4000-m-depth interval is lightly stippled. The locations of four refraction experiments are shown on the map. The ocean bottom hydrophone locations are indicated by the squares and the shooting profiles by the solid lines. The experiment of Detrick and Purdy [1980] is shown by dashed lines; the experiments numbered 2 and 3 are described by Corrnier et al. [1984], as are the results from around the KFZ-MAR intersection on experiment 1. The data presented in this paper are from the line shot down the median valley in experiment 1. The region enclosed by the box is shown in Figure 2. Note that OBH 1 in the 1980 experiment is located almost exactly along a flow line from OBH 1 in this experiment.

considerable and were made using the water path correction method I-Purdy, 1982a]. The seafloor topography as recorded by the conventional wide-beam echo sounder was digitized at closely spaced intervals (generally, < 1 km). These depths were merged with the shot-receiver ranges determined from the direct water wave travel times to produce a depth versus range profile (as shown in Figure 3) independent of navigational uncertainties. The profiles were represented by splines which were used to determine the depth of the ray entry points for any shot and any ray parameter quickly and simply !-Purdy, 1982a]. For display purposes only (i.e., to maintain approximately equispaced seismograms) the record sections shown in this paper have been corrected for topography by removing the water delay time beneath the shot.

Record Sections

Descriptions are given below of the characteristics of the record sections that are important in defining the principal features of the median valley structure. With one exception, only the small-charge (14.5 kg) data are shown because the closer spacing of these seismograms better records the seismic wave field. The large-charge data, with its 3-km shot spacing, was not able to define adequately the important amplitude versus range patterns but was of importance in providing the improved signal to noise ratio necessary to confirm travel time

picks and phase identifications made at longer ranges (35-50 km) on the smaller-charge data.

OBH 1 South. These data (shown in Figure 4a) extend approximately 45 km south of OBH 1, which is situated atop the median valley along-axis high (Figure 3). This is perhaps the most striking record section of all that were collected' it is simple and displays few (if any) unusual features despite being shot along the center of the present-day median valley. It lacks only strong converted shear waves in order to have all the principal characteristics common to most central Atlantic refraction data sets [e.g., Detrick and Purdy, 1980; Purdy, 1983]' low amplitudes at ranges around 15 km due to the low- velocity gradient in layer 3; a high-amplitude Moho triplication caused by a thick well developed Moho transition zone' a weak but clearly visible Pn phase extending beyond 40-km range suggesting a small but positive velocity gradient in the upper mantle; and a strong intercrustal multiple of the Moho triplication delayed by ,-• 1.5 s after the primary. First- arrival travel times could be picked with confidence (to better than + 0.02 s) out to ranges of 45 km.

OBH 1 North. These data extend 80 km north of OBH 1

and are unfortunately marred in the range window 20-30 km by noise generated by a commercial vessel steaming almost directly over the receiving instrument. The most outstanding feature is the lack of identifiable phases at ranges greater than


23o40'N

23o30'N

23o20'N

23o•0'N

23o00'N

22050'N

22o40'N

22o30'N

I I

i I I •

45ø40'W 45ø00'W 44 ø50'W

Fig. 2. Bathymetry contour chart at interval of 200 corrected meters of the MAR median valley and its intersection with the Kane Fracture Zone (at latitude 23ø35'N). The solid contours are from Toomey et al. [1985] except for those around the KFZ-MAR intersection which are from H. J. B. Dick (unpublished data, 1984). Depths shallower than 3000 m are lightly shaded, which approximately de- fines the crestal mountains. The dashed contours are based on wide-

spaced, conventional echo-sounding tracks with general trends biased by visual inspection of the Sea Beam data of Derrick et al. [1984]. The four ocean bottom hydrophone (OBH) positions are shown (see Table 1), and the solid dots denote the explosive charge locations. Major features worthy of note are the nodal deep at the KFZ-MAR intersection that reaches depths in excess of 6000 m, the along-axis topographic high upon which OBH 1 is located, and the apparent disap- pearance of the western crestal mountains north of latitude 23ø10'N. The microearthquake-monitoring experiment described by Toomey et al. [1985] was located within the median valley at approximately latitude 22ø40'N. The bathymetric profile along the explosive shooting line is shown in Figure 3.

I

SOUTH

i i I i i i i

MAR IVlEDIAN VALLEY: ALONG AXIS DEPTH PROFILE

VERTICAL EXAGGERATION 25:4

I

5 400 80 60 40 20 DISTANCE (/cram

I

NORTH

Fig. 3. Bathymetry profile of the axis of the median valley along the small charge shooting track shown in Figure 2. The range coordinate and relative OBH positions were determined from water wave travel time data and thus are independent of navigation. The location of OBH 6 was chosen arbitrarily as the zero range coordinate. Note the location of OBH 1 on top of the along-axis topographic high and the position of latitude 23ø15'N 10 km north of OBH 2. The shaded bands denote possible ridge segment boundaries based on interpretation of gross morphology and magnetic lineations out to anomaly 5 both east and west of the ridge by $chouten et al. [1985b].

40-41 km, where there appears to be a sudden reduction in signal level throughout the seismogram (with the exception, of course, of the water-borne phases). No arrivals are visible at all beyond this point on the 14.5-kg charge record section, and only a severely attenuated first arrival phase can be recognized for a further 10 km on the large charge data. This apparent block in propagation 40 km north of OBH 1 is an important feature, the existence of which is confirmed by the data recorded to the north of OBH 2.

OBH 2 South. These data extend approximately 75 km south of OBH 2, and the 14.5-kg record section (Figure 4b) shows clear first arrivals out to about 47 km range. The 112.6- kg data shown in Figure 4c extend this coverage and display clear first-order and second-order intercrustal multiples, the latter clearly visible at the longest-range shot (~ 76 km). The 14.5-kg data are almost identical to those recorded by OBH ! South out to a range of 25-30 kin, where several significant differences occur. Most importantly, no Pn phase is observable beyond the apparently normal triplication point at 20-25 km but is replaced by a strong coherent lower-velocity phase (~ 7 km/s) that constitutes a clear first arrival on the 14.5-kg data out to 47 km and can be followed on the large-charge data out to about 58 km. The 14.5-kg data also show, unlike those for OBH 1, weak arrivals with travel times consistent with them being identified as a doubly converted shear phase propagating in layer 3. The third important feature of these data is the strong coherent phases arriving within 1 s of the first arrival, particularly on three seismograms centered on 30

TABLE 1. OBH Positions and Depths of the Ocean Bottom Hydrophone Instruments

OBH Depth, Depth,

Latitude Longitude s Two-Way corrected m

22ø54.2'N 44ø57.5'W 4.25 3217 23ø09.7'N 44ø55.1'W 5.23 3958 23ø27.2'N 44ø54.1'W 5.55 4200 23ø32.6'N 44ø52.0'W 6.21 4702


10-

4-

2

0

0

RKAN 01

OBH 1 SOUTH

WATER COLUMN MULTIPLES

2O 3O

RANGE (Iotas)

Fig. 4a. Record section (corrected for topography) of the 14.5-kg (32 lb.) charges detonated south of OBH 1. The seismograms have been truncated at the onset of the direct water wave, and the amplitude scale has been linearly increased with range. The gaps in the section are caused by misfired charges. No filter has been applied. The travel time effects of sea floor topography have been largely removed by subtraction from each seismogram of the water column delay time beneath each shot.

km. These will be interpreted later in the paper: suffice it to say that the first of these, delayed by 0.35 s after the first arrival, is thought to be a peg-leg multiple from a strong reflector in the crust at a depth of approximately 450 m beneath the seafloor.

OBH 2 North. As Figure 4d illustrates, no consistent identifiable arrivals exist on this record section. For purposes of amplitude comparisons these data may be compared directly to those shown in Figure 4b which are data from the same receiver on the same deployment using the same charge type and size. The difference is clear and profound. Beyond 8-10 km range, shots to the north of OBH 2 produce only weak arrivals that are not readily identifiable, although the maxi- mum amplitude arrival does occur at --• 23 km range, consis-

• 4 k.

Fig. 4c.

i i i •10 i I i i K) 40 60 70

t?•NG• (Ams)

As for Figure 4a except that the data shown are the 112.6- kg (248 lb.) charges detonated south of OBH 2.

tent with the peak of the Moho triplication point on both OBH 2 South and OBH 1 South.

This apparent "propagation block" 8-10 km north of OBH 2 coincides with that observed on OBH 1 and the northern-

most small offset transform inferred by Schouten et al. [-1985b] (Figures 2 and 3), but as has been previously mentioned, the multibeam bathymetry data of Derrick et al. [1984] do not support the existence of a transform here. Such severe attenu- ation of energy propagating across a transform fault or fracture zone is also not generally observed [e.g., Sinha and Louden, 1983].

OBH 3 South. This record section displays adequate signal to noise ratio but is complex and difficult to interpret. Two notable features are the post-first-arrival high-amplitude phases occurring between 10 and 15 km range (where they in fact overload the OBH recording system) and at around 25 km. The record section of 112.6-kg charges shows arrivals out to 70 km range, the travel times of which can be picked with some confidence. These data are not immediately interpretable

Fig. 4b.

10

RKAN 01

OBH2 SOUTH

. MULTIPLES• .....

7km/s

i

•0 20 30 40 50

RANGE (Rms)

As for Figure 4a except that the data shown are for the charges detonated south of OBH 2.

2•0 RA•Gœ (:•Os) ' 4• •0 Fig. 4d. As for Figure 4a except that the data shown are the

14.5-kg (32 lb.) charges detonated north of OBH 2. This section should be compared with that shown in Figure 4b. The 20-Hz finback whale vocalizations are highlighted by arrows.


TRAVEL TIME AND AMPLITUDE INTERPRETATION FOR A

lO- • • HORIZONTALLY LAYERED, LATERALLY HOMOGENEOUS MODEL RKAN O1

OBH6 SOUTH Structure to the South of the 23ø15'N Boundary

Explosion travel times. The travel time data used in these

interpretations have been edited to include only those arrivals 6 whose onsets could be identified with confidence to better

- than ___0.02 s. The large- and small-charge data are combined and treated as a single data set. Only first arrivals are con-

4 sidered here, and all have been corrected to the seafloor by using the water path correction method as described previously.

a- . • The data both north and south of OBH 1 are combined in

_ • Figure 5a, and a smooth spline constrained to decrease gradient with increasing range fits the data with an rms deviation of o- 37 ms. Data from ranges greater than 22 km north of the

- instrument have been omitted because of noise interference i i i i i i i ! !

4o ao 3o 4o 5o from a passing commercial vessel and because of the influence RANGE (krns) of the 23ø15'N boundary clearly seen in both the travel times

Fig. 4e. As for Figure 4a except that the data shown are for the charges detonated south of OBH 6.

in terms of simple structures but will be the subject of discussion later in this paper.

OBH 6 South. This OBH, located on the edge of the KFZ- MAR intersection nodal deep, produced incoherent seismograms and complex amplitude patterns (Figure 4e). First- arrival travel times can be reasonably determined out to 55 km from the small (14.5 kg) charge data and to 70 km on the 112.6-kg charge data. These travel times are of considerable interest and are important to the later interpretations, but they should be treated with some caution. As in all seismic interpretations, travel times are one of the most powerful constraints on structures only when the corresponding wave path is known. This is generally inferred from a knowledge of phase velocity, travel time, or amplitude characteristics. However, when heterogeneous structures disrupt commonly identifiable amplitude patterns and a combination of poor spatial sam- pling and scattered travel times result in poor phase velocity determinations, then ambiguity must exist in any interpretation of arrival times alone.

Summary

Two first-order observations may be made based on this qualitative inspection of the seismograms. First, there exists a major structural change (to be called the 23ø15'N boundary) a few kilometers (8-10) north of OBH 2 that severely attenuates all propagation: this is most dramatically illustrated by a comparison of the record sections for OBH 2 South and OBH 2 North (Figures 4b and 4d). Although some energy does cross this boundary, as in the case of the southernmost large charges to OBHs 3 and 6 and the northernmost large charges to OBH 1, no coherent and orderly phases are identifiable. Second, this same boundary apparently separates two distinct- ly different structural regimes: to the south the record sections for OBHs 1 and 2 display coherent phases and consistent amplitude patterns typical of mature oceanic crust. Whereas to the north, OBHs 2, 3, and 6 show attenuated and often incoherent arrivals with no simple amplitude versus range relationships. It is obvious that considerably more detailed determinations of the structure will be possible south of the 23ø15'N boundary than to the north, and thus emphasis will be placed on understanding this southern segment of the median valley.

and amplitudes. The spline curve shown in Figure 5a has little significance other than to affirm the consistency of the travel time data and its compatibility with an interpretation in terms of a horizontally layered, laterally homogeneous structure.

Satisfactory travel time picks could only be made from the shots to the south of OBH 2, and these are shown in Figure 5b: the outstanding feature is the apparent decrease in phase velocity beyond 30 km already obvious from our preliminary inspection of the record section in Figure 4b. As this is clearly not interpretable in terms of a laterally homogeneous structure, it will be dealt with in a later section of this paper: here only the data from ranges less than 30 km will be used. The important observation is that shown in Figure 5c that com- bines the OBH 1 and OBH 2 (<30 km) data but with a constant 0.1 s subtracted from the latter: the two data sets are

seen to be insignificantly different. A spline constrained to decrease gradient with increasing range fits the data (60 points) with an rms deviation of 36 ms. We attributed the 0.1-s offset between the OBH 1 and 2 data to an increased thickness

layer 2A material beneath OBH 2. The impressive similarity in the shape of the travel time curves for OBHs 1 and 2 implies a considerable degree of homogeneity in the structure south of the 23 ø 15'N boundary.

The inversion of the spline curve into a velocity depth function (insert in Figure 5a) provides little more than the starting point for the amplitude modeling but does emphasize the lack of information in this explosion data concerning the uppermost kilometer of the crustal structure. In the structure shown

in Figure 5a this has been approximated by a linear gradient by using the method of Ewing and Purdy [1982]. An independent determination of the structure of the uppermost 2 km will be made based on 16.4-L (1000 cu. in.) air gun data in a later section.

Two fundamental and related characteristics of the struc-

ture are revealed by this simple travel time interpretation: first, the upper mantle delay (intercept) time is 1.5 s, which is approximately 0.5 s greater than that of crust 7 m.y. in age or older [Detrick and Purdy, 1980; Cormier et al., 1984]. Second, the midsection of the crustal column has seismic velocities

significantly lower than are commonly observed (again compared with Cormier et al. [1984] or Detrick and Purdy [1980]): velocities typical of a mature layer 3 (6.8-7.0 km/s) are not observed until a depth of more than 4 km beneath the seafloor (Figure 5a).

Amplitude modeling of explosion data. The amplitude versus range patterns seen in the seismograms recorded from south of the 23ø15'N boundary are simple and are generally

PURDY AND DETRICK: MID-ATLANTIC RIDGE SEISMIC STRUCTURE 3745

2.5

2.0

1.5

1.0

0.5

I I I I I I I " I I

2.5

50

7.5

10.0

0.0

0.0 5.0 10.0 15.0 20.0 25.0 30.0 50.0

RANGE (KMS)

I • I , I , ,

2,5 5.0 7.5 10.0

35.0 40.0 45.0

Fig. 5a. Travel time data for OBH 1 including both 112.6-kg and 14.5-kg shots detonated south of the receiver and those fired north of the receiver at ranges less than 22 km. A spline constrained to decrease its gradient with increasing range is shown that fits the data with an rms deviation of 37 ms. The reduction velocity is 8 km/s. The effects of seafloor topography have been reduced by correcting the shots to the seafloor by using the water path correction method [Purdy, 1982a]. The result of inverting this curve into a velocity depth relationship is shown in the insert. The linear gradient layer above about 1.5-km depth was determined using the method of Ewing and Purdy [1982].

typical of oceanic crust. This can be seen qualitatively in the record sections in Figure 4 but is emphasized by Figure 6. This shows the variation in power in the first 0.25 s of the first refracted arrival on each seismogram versus range for OBHs 1 and 2 South. As in the case of the travel time interpretation,

data from OBH 2 at ranges greater than 30 km are not considered and will be dealt with in a later section of this paper. On OBHs 1 South and 2 South, two clear amplitude peaks at about 11 km and between 20 and 25 km are seen, the latter being due to the Moho triplication. It is these two peaks

2.5

2.0

03 1.5

•--- 1.0

0.5

,

I I I I I I I I I

0.0 5.0 10.0 15.0 20.0 25.0 30.0 35.0 40.0 45.0 50.0

RANGE (KMS)

Fig. 5b. Travel time data for shots detonated south of OBH 2. Note the clear reduction in apparent phase velocity beyond 35 km range. Topographic corrections have been applied as for Figure 5a. Reduction velocity is 8 km/s.


2.5

2.0

L,.i

• 1.0

0.5

I I I I I I I I I

ic m m m m

0.0 5.0 10.0 15.0 20.0 25.6 30.0 35.0 40.0 4.5.0 50.0

RANG•- (KM$) Fig. 5c. Travel time data of Figures 5a and 5b combined but with all data from OBH 2 beyond 30 km range deleted.

In addition, 0.1 s has been subtracted from the OBH 2 data. The spline fits the 60 data points with an rms deviation of 36 ms.

(along with the travel times of course) that are the primary features to be modeled here. The similarity in the amplitude- range i'elationships of OBHs 1 and 2 is striking and provides further justification for making a first-order interpretation of

1.0 I ' I I I

0.8

0.4

0.2

0.0 0.0 10.0 20.0 30.0 40.0 50.0

RANGE (KMS) Fig. 6. Normalized "power" versus range for the 14.5-kg charges

north (pluses) and south (crosses) of OBH 1 and south (stars) of OBH 2. The "power" is determined simply by summing the squares of the amplitudes in the first 0.25 s of the first refracted arrival. The values have been scale& up by the square of the range. Note the good agreement between OBH 1 South and OBH 2 South out to 30 km. Beyond this the high-amplitude 7 km/s apparent velocity observed on OBH 2 results in significant departure from the OBH 1 data. The single, clear Moho triplication high observed on OBH 1 South and OBH 2 South is not observed on OBH 1 North.

these data in terms of a horizontally layered, laterally homogeneous structure.

The WKBJ synthetic seismograms that satisfactorily match the observed data are shown in Figure 7. The travel time curve shown overlain on the combined OBH 1 travel times in

Figure 8a are calculated from the velocity-depth structure shown in Figure 8b and listed in Table 2. This is our best estimate for the structure of the crust and upper mantle south of the 23ø15'N boundary in terms of a horizontally layered, laterally homogeneous model based on the explosive amplitudes and travel times recorded south of OBH 1, to the horth, of OBH 1 at ranges less than 22 km, and to the south of OBH 2 at ranges less than 30 km.

The important features to be modeled are the two major amplitude peaks at 11 km and between 20 and 25 km, the rate of decay of the upper mantle refraction, and the lack of wide- angle reflections from the Moho (PtnP). Referring to Table 2 and Figure 8b, the shallowmost layer with its seafloor velocity of 2.4 km/s and gradient of 2.45 s-• is required only to match the travel times and is not well-constrained. This single layer is intended to represent a mean estimate of the shallow structure over the • 80-km length of the median valley sampled by these data. On a smaller scale the structure of the uppermost i.1 km of the crust is more complex and variable (as will be described later). However, for the purposes of this large-scale model of the mean structure south of the 23ø15'N boundary, this single, linear, gradient layer is judged to be a satisfactory approximation. The lack of observed converted shear waves supports the existence of a low seafloor velocity that will produce an inefficient boundary for the conversion of compressional to shear wave energy (or vice versa).

The amplitude high at 11 km range is a result of the focusing of energy by the 1.4-km-thick layer of gradient 0.71 s -• that underlies this shallowmost layer. The velocity at the base of this layer is only 6.1 km/s, although it is at a depth of 2.5 km beneath the seafloor. A considerable reduction in velocity


6.0

5.5

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I I I I I I I

OBS R

I I I I I

0.0 5.0 10.0 15.0 20.0 25.0 ,30.0 ,35.0 40.0 45.0 50.0

6.0

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I I I I

SY qTHETIC

I ,35.0 40.0 45.0

,3.0 I , I I I I 0.0 5.0 10.0 15.0 20.0 25.0 ,30.0

RANGE (KMS) Fig. 7. Observed and WKBJ synthetic seismograms for the 14.5-kg charges south of OBH 1. The travel time curve

and velocity depth function for the best fit model are shown in Figure 8. The observed seismograms at approximately 26 and 27.5 km range have very weak first arrivals: the bubble pulse periods (recorded by the towed hydrophone used to monitor the shot instants) were anomalously high for both these shots, suggesting a partial detonation. Because of this, no significance is attached to the fact that the synthetic Moho triplication is apparently broader than the observed. Reduction velocity for both sections is 8 km/s. For this figure the observed data have been corrected to a datum model with a constant 4-km thickness water layer.

50.0

gradient is needed below this layer in order to model both the sharp decrease in amplitudes to a minimum around 15-18 km and the lack of curvature in the observed travel time relation-

ships until 20 km range. This 0.19 s-• gradient layer (layer "Y'

in Figure 8b and Table 2) can only be --,0.8 km thick because of the need to build up amplitude at 30 km range and begin the increase in phase velocity required by the travel times. Layer "4" serves to increase the phase velocity and contribute

3748 PURDY AND DETRICK' MID-ATLANTIC RIDGE SEISMIC STRUCTURE

2.0

1.5

1.0

0.5

+ -

_

0.0 5.0 .10.0 '15.0 20.0 25.0 30.0 55.0 40.0 45.0

RANGE (KMS) Fig. 8a. Travel time curve computed for the velocity depth func-

tion shown in Figure 8b superimposed on the OBH 1 travel time data. Reduction velocity is 8 km/s, and the data have been water path corrected to the seafloor.

to the principal amplitude high, and layer "5" is required to slow the rate of curvature of the travel time curve and broa-

den the amplitude high, bringing the velocity up to 6.85 km/s. This is inferred to be the base of the crust, the Moho transition itself consisting of the 2.2-km-thick combination of layers 6 and 7 that overlay the 7.9 km/s, 0.02 s-x gradient upper mantle. This thickness of the Moho transition results in little

or no Prop energy (thus matching the observations), and the comparatively large upper mantle velocity gradient adequately

0.0

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4.0

5.0

6.0

7.0

8.0

2.0 3.0 4.0 5.0 6.0 7.0 8.0 9.0

VELOCITY (krns/sec)

Fig. 8b. Velocity-depth relationship for the best fit solution for OBH 1. The travel time curve and WKBJ synthetic seismograms corresponding to this structure are shown in Figures 8a and 7, respectively. The depth is relative to the seafloor. The layer numbers are referred to in the text. Velocities and thicknesses are given in Table 2.

TABLE 2. Mean Structure South of the 23ø15'N Boundary

Velocity, km/s Thickness, Thickness,

Layer km Top Bottom s Two-Way

1 1.1 2.40 5.10 0.614 2 1.4 5.10 6.10 0.501 3 0.8 6.10 6.25 0.259 4 1.2 6.25 6.80 0.368 5 0.5 6.80 6.85 0.146 6 1.2 6.85 7.60 0.332 7 1.0 7.60 7.90 0.258 8 3.0 7.90 7.95

models the strong Pn refractions. Relative thicknesses and relative velocity gradients of the layers 6 and 7 that constitute the Moho are not well-constrained.

Shallow structure from air •lun data. Although air gun profiles were carried out over all four OBH instruments, useful data pertaining to the structure of the shallow crust were obtained only by OBH 1. A record section of these data south of the instrument is shown in Figure 9. A strong low-velocity phase is seen to emerge tangentially from the water wave at • 3.8 km range and continue out to almost 7 km. When the travel times of this phase are corrected to the seafloor, the intercept time, within error, is zero. This is an unusual observation [cf. Derrick and Purdy, 1980; Ewing and Purdy, 1982]. The zero intercept time and tangential relationship with the water wave means that this phase (if it is indeed a refraction) must be propagating at or close to the seafloor. The second phase that dominates the record section is of much higher

OBSERVED •5 -

• •o

2.5

2.0 I 2 ..... 0

4.0

2.5

2.0

2.0 • 0 4.0 5 0 6.0 7.0 8 0

RANGE (KMS)

Fig. 9. Observed and WKBJ synthetic seismograms for the 1000 cu. in. air gun line south of OB• 1. The velocity depth function corresponding to the best fit mode] is shown in Figure 10. The observed data have been truncated at the onset of the direct water wave.

This emphasizes the emergence at about 4 km range of the • 2.5 km/s phase propagating in a thin layer below the seafloor. Reduction velocity is 8 km/s.


0.00

0.25

0.50

0.75

1.00

1.25

1.50

1.75

2.00

2.25

2.50

2.75

AIR GUN

3.00 , I , I , I , I / , 2.0 3.0 4.0 5.0 6.0

VELOCITY (kms/sec) 7.0

Fig. 10. Velocity depth function for the best fit solution for the 1000 cu. in. air gun data south of OBH 1 shown in Figure 9. Veloci- ties and thicknesses are given in Table 3.

phase velocity, is extremely coherent, and is focused into a 1-km-wide range window centered on 6 km range.

A synthetic seismogram section that matches the observed record section satisfactorily is shown in Figure 9. The velocity- depth function from which the synthetic seismograms in Figure 9 were calculated is shown in Figure 10 and Table 3. The low gradient shallow layer (0.22 s-•) is required to maintain the low-velocity "seafloor" phase out to •7 km range. This 450-m-thick layer is underlain by a first-order discontinuity over a 3.9 km/s layer. We cannot resolve whether the 2.6 km/s phase that emerges from the water wave at •4 km is a refraction propagating in this 450-m-thick low-gradient layer or is a wide-angle reflection from a discontinuity at its base. The travel time difference between these phases is in fact insig- nificant, and our preliminary amplitude analyses lack the rigor to make the distinction important. Whether it is a reflection or a refraction, a low-gradient layer is needed to maintain the amplitudes out to ~ 7 km. The isovelocity layer is needed to reduce the amplitude at around 5 km range, but the velocity of 3.9 km/s is poorly constrained. The 1.1-km-thick transition zone that with a gradient of 2.2 s-• brings the velocity up to 6.1 km/s is required to produce the focusing of coherent energy at ~6 km range. This is a well-determined feature of the velocity-depth section and is most probably a universal characteristic of young oceanic crustal structure [e.g., Bratt and Purdy, 1984]. However, the low-gradient layer immediately beneath the seafloor is unusual: Although such low velocities have been previously suggested [e.g., Whitmarsh, 1978; Ewing and Purdy, 1982], high gradients have been required to satisfy the observed travel time and amplitude relationships.

There seems little doubt such a low-velocity, low-gradient layer is needed to satisfy our observations: No other satisfactory explanation for this phase could be found, and it was clearly observed in the data both to the north and south of OBH 1. However, despite exhaustive trials it was not found to be possible to match the travel time and amplitude patterns of the explosive data described in the previous section using the shallow structure determined from the air gun data. The infer- ence is that this structure is restricted to a small region (• 5 km radius) around OBH 1 beneath the peak of the along-axis topographic high. Independent supporting evidence for this will be presented later.

Structure to the North of the 23ø15'N Boundary

In contrast to the previous section, little data are available to determine the structure north of the 23ø15'N boundary. Only the travel time data for OBH 6 produce a useful solution. The complexity of the OBH 3 record section and the larger uncertainties in the first arrival time picks make any solution for the instrument of doubtful value.

The primary characteristics of the OBH 6 travel time data set have been previously described by Cormier et al. [1984]: They are the clear reduction in apparent phase velocity from ~ 8 km/s to ~ 7 km/s at 30 km range (see Figure 13 of Cor- mier et al. [1984] and later in this paper) and the fact that the observed travel times are as much as 0.5 s less than those for

OBH 1 (cf. Figures 5a and 11 or also see Figure 13 of Cormlet et al. [1984]). The reduction in apparent phase velocity at 30 km range naturally restricts any interpretation of the data in terms of a laterally homogeneous, horizontally layered structure to ranges less than this. The travel times used for this solution along with the preferred model travel time curve is shown in Figure 11. The corresponding velocity depth function is included in Figure 12b. Because of the lack of coherent amplitude patterns that can be satisfactorily modeled using synthetic seismograms, the details of the layering of this model are not well-constrained. However, the depth below the seafloor at which upper mantle velocities are observed is 2-3 km less than that determined for OBHs 1 and 2: That this differ-

ence is real and significant is clear from the previously described comparison of the observed travel times.

Although it is not observed as clearly on OBH 3, apparently the same reduction in phase velocity that is seen at • 30 km range on OBH 6 is seen at •20 km range on OBH 3. This reduced range of available data makes a determination of a laterally homogeneous solution of little value. The important observation, however, is that the reduction in apparent phase velocity observed on OBHs 6 and 3 occurs at the same geo- graphic location (the two receivers are separated by • 10 km) and thus is presumably a consequence of a major change in structure beneath the shot positions. This is convincingly illustrated by Figure 13, that simply compares the observed travel times for OBHs 6 and 3 but with the ranges for all the OBH 3 observations increased by 10 km.

TABLE 3. Structure Beneath OBH 1 From Air Gun Data

Velocity, km/s Thickness, Thickness,

Layer km Top Bottom s Two-Way

I 0.45 2.55 2.65 0.34 2 0.50 3.90 3.90 0.26 3 1.10 3.90 6.10 0.45


2.0

1.5

1.0

0.5

+

0.0 ,5.0 10.0 1,5.0 20.0 25.0 ,30.0 ,35.0 40.0 45.0

RANGE (KMS) Fig. 11. Observed travel time data from 14.5-kg charges detonated south of OBH 6 along with travel time curves

corresponding to one possible structural solution shown in Figure 12b. Delayed arrivals beyond ,--30 km range cannot be accounted for in terms of a laterally homogeneous model. The observed travel time data have been corrected to the seafloor, and the reduction velocity is 8 km/s. The record section from which the travel time picks were made is shown in Figure 4e.

Summary of Homogeneous Structure Determination

A number of important facts have been established that are worthy of review before beginning construction of the two- dimensional model. The 100-km section of ridge that is the subject of this study is characterized by two significantly different crustal structures: OBHs 1 and 2 reveal "normal" crust-

al thickness, a considerable degree of lateral homogeneity but unusually low velocities (cf. mature oceanic crust) particularly in the upper part of layer 3 (Figure 12); and OBHs 6 and 3 reveal thinner crust (Figure 12) that is inhomogeneous on lateral scales of several kilometers, as is suggested by the lack of coherent and systematic amplitude-range relationships.

Perhaps the singlemost significant result is the determination of 7.9-8.0 km/s upper mantle velocities beneath the median valley. Although as was described in the introduction, other data sets have suggested the presence of "normal" upper mantle beneath the median valley [e.g., Poehls, 1974; Whit- marsh, 1975], the observation presented here is unequivocal. To support this statement, Figure 14 shows the seismograms from shots located 28-42 km south of OBH 1 at an expanded normalized amplitude scale and a compressed range scale (compared with that shown in Figure 4a). This establishes the clarity and coherency of this first ariving Pn phase. It shows that the phase velocity determination is not a fortuitous consequence of random errors in picking times of first onset but is truly the phase velocity of the first arriving coherent wave packet. The laterally homogeneous, horizontally layered structures determined so far in this paper will now be used as the primary building blocks of a more complex two- dimensional model of the along-axis changes in structure beneath the median valley.

A Two-DIMENSIoNAL MODEL OF ALONG-AXIS CHANGES IN STRUCTURE BENEATH THE MEDIAN VALLEY

Although the previously described structural solutions account for most of the primary travel time and amplitude characteristics of the record sections shown in Figure 4, five important features remain unexplained:

1. At ranges greater than approximately 30 km south of OBH 2, a clear and coherent ,-• 7 km/s phase becomes the first arrival (Figures 4b and 5b). This occurs following a Moho triplication point that is, in terms of range, width, and power

level, insignificantly different from that of OBH 1 (as illustrated by Figure 6). The range window over which this phase is observed on the small-charge data approximately coincides with the 1.8 ø along-axis slope in the seafloor south of OBH 1 (compare Figures 4b or 5b with Figure 3). Thus explanations for this phase were sought initially in terms of topographic effects. The observed phase velocity uncorrected for topography was approximately 6 km/s. Water path corrections for topography FPurdy, 1982a] increase this to ,-• 7 km/s. Nonran- dom errors in water depth at the ray entry points of up to several hundred meters would be necessary to change the topographic corrections sufficiently to bring this phase velocity up to 7.9 km/s. This is considered to be an unlikely explanation because of the large number of along-axis echo- sounding tracks collected during this experiment for the three coincident shooting lines (as well as lines for the instrument deployments and recoveries). Second, the same water depth profile was used to correct the shorter range data south of OBH 1: the proof that the topographic corrections for these are reliable is the excellent agreement between the travel times north and south of OBH 1 at ranges less than ,-•25 km (Figure 5a). The third piece of evidence that this phase is not an artifact of the topographic corrections is the straightfor- ward observation that it extends south of the southern limit of

the 1.8 ø slope (approximately at the 90-km coordinate in Figure 3). This is not clear from the small-charge data shown in Figures 4b and 5b, but the large-charge data carry this same low-phase velocity out to ,-• 58-km range, i.e., to the ,-, 100 km coordinate in Figure 3 or 10 km beyond the southern limit of the 1.8 ø slope. Figure 4c shows a prominent reduction in the amplitude of this phase at 48 km: Whether or not this amplitude decrease is the consequence of some deep structure or whether it is a topographic effect must remain a matter of some conjecture.

The clarity and coherency of the arrivals, our inability to cast doubt on the validity of the topographic corrections, and the presence of the phase on both the large and small data sets mean that this 7 km/s phase must be treated as a well- constrained observation. Any structural model for this section of the median valley should provide an explanation of this phase.

2. Some kind of major structural boundary must exist


0.0

1.0

2.0

3.0

4.0

5.0

6.0

7.0

\

\

\

\

\

OBH 1 -->

80 , I , I , I , I • 2.0 3.0 4.0 5.0 6.0

VELOCITY (kms/sec)

<-- 7 my

•<-- OBH 6

ß

,,

i

1

1

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Fig. 12. Solid line is one possible (poorly constrained) velocity depth structure for the OBH 6 data, the travel time curve for which is illustrated in Figure 11. For comparison, the dashed line shows the OBH 1 solution as previously illustrated in Figure 8b. The difference between these solutions that has significance is the depth below the seafloor at which upper mantle velocities is reached: 6-7 km on OBH 1 compared with -•5 km on OBH 6. The OBH 1 solution represents the structure south of the 23ø15'N boundary, and the OBH 6 structure is representative of that to the north. These are compared with the solution from Detrick and Purdy [-1980] for the shots north of OBH 1 (in the 1980 experiment) located along a flow line of 7-m.y.-old crust (see location in Figure 1).

--• 10 km north of OBH 2 that blocks, or at least severely

attenuates, propagation at all levels through the crust. The clearest observation of the effects of this boundary at latitude 23ø15'N is the comparison of the record sections south and north of OBH 2 (Figures 4b and 4d). Its effects are also clearly seen on OBH 1, where no identifiable arrivals exist to the north at ranges greater than 40 km. As has been previously mentioned, the OBH 3 record section is extremely complex, and although large changes in amplitude occur in the range window 20-30 km, it is not clear that they are related to this same phenomenon. The case is somewhat similar with OBH 6 (Figure 4e): The arrivals are relatively weak and incoherent everywhere, so relating some particular amplitude change to this presumed structural discontinuity (which for OBH 6 would occur at 30 km range) is not possible.

3. The first-arrival travel times to both OBHs 3 and 6,

though not of great help in constraining the homogeneous solutions because of their low signal to noise ratio and incoherent waveforms, produce a remarkably consistent observation in that they both seem to indicate a reduction in phase velocity at ranges of 20 and 30 km, respectively. The validity

and consistency of this observation is well-established by Figure 13. It is presumably not a coincidence that the location of this reduction in phase velocity is that of the propagation block discussed above. A qualitative interpretation would be that the crustal thickening required to change the OBH 6 structure to the OBH 1 structure (Figure 12) occurs at this location and is the cause of the lowered phase velocities.

4. Although the loss of signal 40 km north of OBH 1 has been discussed, a second important characteristic of the OBH 1 North record section exists at 20-25 km range. If the structure south of the 23ø15'N boundary were as simple and homogeneous as some of our results seem to suggest, then we should expect excellent consistency of the amplitude and travel time patterns observed on OBH 1 South. This is indeed the case out to ,-•22 km range: to this point the travel times for the two data sets are indistinguishable, but beyond this range the OBH 1 North travel times are delayed relative to those on OBH 1 South. In addition, although the amplitudes begin increasing on OBH 1 North at 15-20 km, as they do on both OBHs 1 South and 2 South (Figure 6a), a single broad amplitude high corresponding to the Moho triplication point


2.0

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RANGE (KMS) Fig. 13. A comparison of the water path corrected first-arrival

travel times from charges detonated to the south of OBHs 6 and 3. The range of all the OBH 3 data has been increased by 10 km (which is the separation of the two receivers). The data do not coincide in range precisely because of the differing range shifts associated with the water path corrections and the fact that OBH 6 lies several kilometers east of the shooting line (see Figure 2). As the data points are plotted with a reduction velocity of 8 km/s, an alternative view of this figure is one of total crustal delay times plotted versus range to OBH 6. The important point is the excellent agreement between the two data sets in the location of the reduction in apparent phase velocity (• 30 km). This proves this must be due to an along-axis change in structure occurring beneath the shot points.

is not observed, but rather two narrower peaks of lesser amplitude. An explanation for this must be sought by ray tracing.

5. Three seismograms recorded on OBH 2 South display a phase of particular interest (Figure 15). These seismograms (at

range of 28-31 km) are located over OBH 1 and the peak of the along-axis topographic high. They each display three coherent arrivals, the second of which is phase-shifted by rc relative to the first (this is especially clear on shot 162 shown in Figure 15). The separation of these first two phases is 0.3 s on shot 161 and 0.35 s on shots 162 and 164. The interpretation of the air gun data recorded by OBH 1 produced a velocity- depth function with an abrupt 1.25 km/s discontinuity 0.34-s two-way reflection time beneath the seafloor. The similarity between the thickness of this surficial layer and the separation of these first two phases, the presence of a strong discontinuity at the base of this layer, and the phase shift of r• of the second arrival relative to the first support the identification of this second arrival as a peg-leg multiple in the low-velocity surficial layer beneath the shots. The phase shift results from the near-normal (• 12 ø for a 7 km/s phase velocity) incidence reflection beneath the seafloor. It has already been suggested that the shallow structure for OBH 1 as determined by the air gun data is appropriate only to a small area around the peak of the along-axis topographic high. The presence of this peg- leg multiple on only those shots within a few kilometers of OBH 1 supports this tentative conclusion. If this interpretation is correct, then this same phase should also be recog- nizable on OBH 1: in this case, of course, the peg-leg reflection would be taking place beneath the receiver and not beneath the shot. Although it lacks the clarity of Figure 15, this same phase can indeed be recognized on OBH 1 (Figure 14). The separation is 0.3 s, and it is phase-shifted by r•.

Unfortunately, no such consistent explanation can be found for the third phase observed in Figure 15. Some interference distorts and lengthens the waveform of this phase in shots 161

lO o

9.5

9.0

85

8.0

7.5

7.0

Fig. 14. This figure illustrates the coherency of the Pn phase observed on the 14.5-kg charges in the range window 28-42 detonated south of OBH 1 and thus emphasizes the confidence with which normal upper mantle velocities are determined beneath the median valley. No filtering has been applied. Each seismogram has been separately amplitude scaled. Compression is to the left. The bold chain dashed line delineates the refracted Pn phase. The faint dots denote a possible peg leg multiple in a thin layer beneath the OBH: see text and Figure 15 for further discussion.


Fig. 15.

9.0

8.5

8.0

7.5

7.0

Seismograms for shots 161, 162, and 164 at ranges of 28.2, 29.8 and 31.4 km, respectively, south of OBH 2 (shot 163 misfired). Note that OBH 1 and the peak of the along-axis topographic high was located 29.5 km south of OBH 2 (see Figure 3). We interpret the coherent secondary phase to be phase-shifted by n relative to the first arrival:this is most clearly seen on shot 162. The separation of these first two phases is then 0.35 s, identical with the thickness of the shallow low-velocity layer determined from the air gun data south of OBH 1 (see Table 3 and Figure 10). The secondary phase is thus interpreted to be a peg-leg multiple within this shallow layer.

and 164, and it is not observed on OBH 1 (Figure 14). No reliable identification of this phase can be made.

Construction of a Two-Dimensional Model

The structural model used for the ray tracing that satisfactorily accounts for most of the features observed in the data is defined in Figures 16-18. A set of analytical functions is used to define the velocity field through which the rays will be traced. The foremost of these is the spline that approximates the along-axis bathymetry profile. Identical splines, simply shifted in depth, are used to describe the top of layer 3 and the top of the M oho. The velocity depth functions are described by the exponential integral shown in Figure 16: three sets of parameters describe the complete crustal column, one set each for the upper crust, layer 3, and the Moho and upper mantle. All these parameters are included in Figure 16. The velocity- depth functions described by these parameters are shown in Figure 17 along with the linear gradient layer solutions resulting from the laterally homogeneous interpretations. The model has been divided into five lateral zones each of which

has its own velocity-depth function. The splines define the zero for the exponential integrals, and thus, as is obvious from Figure 18, the isovelocity contours are everywhere conformal with the seafloor (with the exception of one small area > 10 km depth in zone 4). The zone boundaries are simply vertical, first-order discontinuities.

The model is simple, only minor modifications being made to the structures determined under the assumptions of lateral homogeneity. It is difficult to know how best to construct a two-dimensional model that will satisfactorily account for the observations and have a reasonable degree of uniqueness. The ambiguity of two-dimensional inhomogeneous structural solutions that are constrained only by ray-traced travel times is well-known. There is no doubt that by constructing a complex

model with many small-scale structural inhomogeneities, the observed travel times could be matched more closely than will be the case here. Such a model would simply be one of an infinite number of widely varying solutions that could fit the observations satisfactorily. The approach taken here is direc- ted toward retaining as far as possible the uniqueness of the laterally homogeneous solutions. The question posed while building the model was: How well can the five unexplained features of the data described in the previous paragraphs be accounted for whilst making the minimum number of changes to the structures determined from travel time inversion and

amplitude modelling?. In this way the resulting model will be far less ambiguous than a more complex model that accounted for every detail of the observations.

The second point to be made is that the model shown in Figure 18 is a model that explains the seismic observations and is not a geological model. The latter will no doubt change and evolve as our understanding of the processes improves; the former will, providing no errors exist here, stand as fact, to which time will bring only detailed enhancements.

In constructing the model described here only two departures from the laterally homogeneous structures were required in addition to three "semiarbitrary" decisions, the first of which was to require the isovelocity contours to be conformal with the topography. This seems not only a physically more reasonable situation than forcing horizontal layering (especially in the shallow crust [Purdy, 1982b], but it is also more compatible with the water path topography correction approach [Purdy, 1982a]). The remaining two decisions were concerned with the location of the known along-axis structural changes. The largest change is that from OBH 6 structure to OBH 2 structure: for obvious reasons, that was located at the 23ø15'N discontinuity 10 km north of OBH 2, i.e., at coordinate 30 km in Figures 16 and 18. The second lesser structural boundary between OBH 1 and OBH 2 type crust was located at the 50-km coordinate. The reason for this choice will

become clear during the presentation of the ray-tracing results. The two departures from the laterally homogeneous structures were the introduction of a 15-km-wide zone of low-

ered velocities beneath OBH 1 (i.e., zone 4) and, second, the reduction in upper mantle velocities by 0.3-0.4 km/s below OBH 2. Both these modifications will be justified by the ray trace calculations.

Ray- Tracinft Results

Examples of ray paths through the structure described above are shown in Figures 19 and 20. Comparisons of the observed travel times with those computed from the ray tracing are shown in Figures 21 and 22. The ray trace calculations were carried out as described by Purdy [1982b], with the added allowance for refraction at the zone boundaries. With a

model as large as that shown in Figure 18, the advantage of using a set of analytical functions to describe the velocity field quickly became clear. As is shown in Figure 16, apart from the splines, only an additional 33 parameters are needed to define uniquely the velocity distribution that is contoured in Figure 18.

The ray paths south of OBH 1 shown in Figure 19 represent the simplest case in the model, as no zone boundary crossings or complex structures are involved: they are almost exclusively restricted to zone 5, the velocity-depth function for which is shown in Figure 17. Rays are seen to be focused at •-- 10 km range (corresponding to the similarly located power peak shown in Figure 6) and at 17 km (top of layer 3) and 23 km (top of Moho). When the effects of finite wavelength are accounted for (as they are, of course, in the computation of the


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6

8

Fig. 17. The velocity-depth relations in each of the five along-axis zones defined in Figure 16. In each case a linear gradient model is shown that was determined from travel time inversions and synthetic seismogram modeling, overlain by a three-layer, exponential function model that was used in the ray trace model. This function and all the variables are defined in Figure 16. See text for further explanation.

synthetic seismograms shown in Figure 7), these latter two concentrations of energy merge to form a single broader peak. But this is exactly as it would be in the simple laterally homogeneous case: for the OBH 1 South data this two-dimensional ray tracing does little other than confirm the consistency of our calculations. The match between the observed and the

calculated travel times is shown in Figure 21. Beyond approximately 25 km range the calculated travel times are seen to be delayed, on average, by a few hundredths of a second compared with the observations. This is due to the "apparent" thickening of the crust sensed by the rays propagating through this laterally inhomogeneous structure, simply caused by the fact that OBH 1 is located on a topographic high.

As was described earlier, one of the subjective decisions required in building the model was the location of the small change from OBH 1 to OBH 2-type structure. If this boundary is positioned at the 50-km distance coordinate, then it provides a possible explanation for the retarded travel times on OBH 1 North (compared with those of OBH 1 South) at ranges greater than --• 25 km and for the distorted amplitude pattern of the Moho triplication point (Figure 6). No attempt is made to model this latter effect: attention is drawn only to the interesting coincidence in location. Of course, in reality it is most unlikely that the boundary between the OBH 1 and 2 structures is a discrete vertical first-order discontinuity, but as was emphasized earlier, this is a seismic model, and the existing seismic data are incapable of resolving the nature of the boundary. Thus the choice of a vertical discontinuity is because it is simple, and the effects are readily understood and predicted. For example, the details of the distortion in the amplitude relationships of the Moho triplication point on OBH 1 North may be explainable in terms of the precise character of the lateral transition at this boundary. Such detailed and ill-constrained modeling is beyond the capability of

• 4

0 I I I I I I I I ! I ! I I I I I I i I I I I

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Fig. 18. An isovelocity contour representation (contour interval mostly 0.5 km/s) of the structure used in the ray tracing and previously defined by Figures 16 and 17. Note the anomalous zone of low velocities beneath the along-axis topographic high at the 70-km range coordinate and the large and sudden decrease in crustal thickness at the 30-km coordinate. The encircled numbers refer to the zones defined in Figures 16 and 17.


2 -- SOUTH RAY PATHS NORTH AND SOUTH OF OBH '1 VERTICAL EXAGGERATION '-.'5.'• NORTH - OBH "1 OBH 2

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Fig. 19. Rays traced to the north and south of OBH 1. The dotted lines are those of Figure 16. Note how the anomalous zone beneath OBH 1 has no significant effect upon propagation to OBH 1. See text for discussion.

the methods employed here as well as being beyond the scope of this paper. The single most important feature of the ray diagram shown in Figure 19 is how little, if at all, the propagation to OBH 1 is affected by the zone 4 structure centered beneath the peak of the along-axis topographic high.

In contrast with the energy propagating from south of OBH 1, the rays to the south of OBH 2 are profoundly effected by lateral changes in structure. Three important features of the OBH 2 data set were (1) the •,0.1-s delay of all travel times compared with those of OBH 1 (Figure 5c); with the OBH 1 to OBH 2 structure boundary located at distance coordinate 50 km, the rays arriving at OBH 2 propagate through the complete section of thickened shallow crust beneath OBH 2 and thus become delayed, (2) the similarity in range, width, and level of the Moho triplication with that of OBH 1 South (Figure 6); the rays that make up the Moho triplication point for OBH 2 South all bottomed out between distance coordi-

nates 50 and 63 km (zone 3), which is a section of OBH 1-type crust, and thus the resulting amplitude.patterns should be similar to those observed on OBH 1 South, (3) last, and perhaps most importantly, the unusual ,-,7.0 km/s phase observed most clearly in the range window 30-48 km. This is explained by the otherwise unconstrained addition of a 15- km-wide zone of lowered velocities beneath the peak of the along-axis topographic high. As can be seen in Figure 20, energy that would have been turned sufficiently to contribute to the triplication point in the laterally homogeneous case crosses the boundary into zone 4, experiences the lower gradients and velocities that exist there, and so is refracted out to longer ranges. The delays caused by this longer path and the passage through the lower-velocity material result in the 7 km/s apparent phase velocity. The match between the observed and the calculated travel times is shown in Figure 21. It

is important to refer to the zone • velocity-depth function in Figure 17 to understand how this structural model is re- distributing the energy: there is no velocity inversion within the "crust," only a significantly lowered velocity and velocity gradient. This takes the energy that would have been focused into the Moho triplication point and spreads it out over the range window 30-48 km. The ve19city inversion that has been placed at 10 km below sea level is required to block long- range, upper mantle propagation and to cause a major decrease in the amplitude of the 7 km/s phase at ,-,48 km (in agreement with the observations in Figure 4c). No meaningful constraints exist on this feature of the model: a highly attenu- ative zone or an infinite number of alternative velocity structures could produce this same effect.

Models involving a 10- to 15-km-wide zone of reduced velocity in the lower crust beneath the peak of the along-axis topographic high were the only ones that could be found that produced a reasonable fit to the data. However, the model presented here does not account for the amplitude versus range relationship of the 7 km/s phase beyond 30 km in one important way. Following the triplication point, amplitude high at 20-25 km range, there seems to be a window of lowered amplitudes around 30 km range before the amplitudes increase out to ,-, 40 km range (Figure 6). The model presented here does not predict this decrease in energy of the first arrival around 30 km. There are three possible explanations for this: (1) It is a consequence of defocusing of energy through the water column by the along-axis topographic high (the peak of which lies at 30 km range from OBH 2), (2) it is a consequence of the highly reflective discontinuity in the shallow crust that was observed on the OBH 1 air gun data and, as described earlier, is thought to generate the peg-leg multiples seen on these same shots to OBH 1 around 30 km range (Figure 15),

PURDY AND DETRICK' MID-ATLANTIC RIDGE SEISMIC STRUCTURE '3757

.12

SOUTH RAY PATHS SOUTH OF OBH 2 AND OBH 6 VERTICAL EXAGGERATION ,',, 5 '1 NORTH

OBH -1

,

'140 .100 90 80 70 60 50

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OBH 2

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Fig. 20. Rays traced to the south of OBHs 2 and 6. Note how the anomalous zone beneath OBH 1 significantly effects propagation to OBH 2 beyond about 30 km range. See text for discussion.

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and (3) it is a consequence of lateral changes in velocity structure within the lower crustal zone of reduced velocity. By inspection of Figure 20 it can be seen that if the velocity at the northern boundary of zone 4 was lowered significantly below its present value, a shadow zone at --• 30 km could be created due to the refraction (toward the normal) of the upcoming energy at the zone boundary. A lateral increase in velocity toward the south could be used to focus energy at •40 km followed perhaps by a second decrease at the southernmost zone boundary to contribute toward the drop in amplitude at 48 km range.

Of course, all three of the above could be contributing fac- tors, and an assessment of their relative significance requires the calculation of synthetic seismograms through a laterally inhomogeneous medium. These data are worthy of such a detailed study, but it lies beyond the scope of this paper.

The ray paths south of OBH 6 are dominated by the effects of the crustal thickening at the 23ø15'N boundary (between zones 1 and 2) located 10 km north of OBH 2. Predictably, this causes a significant delay in lowering of the apparent phase velocity that is clearly observed in the data (Figure 13) and in the calculated travel times (Figure 22). This confirms the coincidence in location of the change in thickness from OBH 6 to OBH 2 type crust with that of the propagation block referred to as the 23ø15'N boundary. However, models involving simply the crustal thickening do not satisfactorily fit the observed travel times: they produce the decrease in phase velocity at 30 km (as is observed), but this extends over a range of only • 10 km: propagation then returns to upper mantle phase velocities though, of course, delayed due to the passage through the thicker section. The observed travel times clearly show decreased phase velocities out to at least 50 km range:to reproduce this in the model, the upper mantle veloc-

ity within zone 2 has been arbitrarily reduced. The isovelocity contour representation of the model tends to exaggerate the significance of this adjustment: a more appropriate view is that of Figure 17. This feature is without question the most ill-constrained characteristic of the model. The limits of this

lowered upper mantle velocity were chosen to be the 23ø15'N boundary and the OBH 1/2 structure boundary, but this need not be the case. The southern extent cannot be much greater than the •55 km coordinate because, otherwise, its effects would be observed in the OBH 2 South triplication point. The northern limit is constrained only by scattered OBH 3 travel times.

Summary of Ray- Tracing Results

This modeling of the observed travel times by ray tracing through a laterally inhomogeneous structure establishes two important facts that could not be determined from the travel time inversion and synthetic seismogram modeling using horizontally layered, laterally homogeneous, velocity-depth functions. First, it proves the coincidence in location of the abrupt change from "OBH 1-type" or "thick" oceanic crust to "OBH 6-type" or "thinner" crust with both the major propagation block observed on OBH 2 and the truncation of the western

crestal mountains just south of latitude 23ø15'N. Second, the ray tracing shows that a 10- to 15-km-wide zone of significantly lowered (by 0.5-1.0 kin/s) velocities must exist at the base of the crust beneath the along-axis topographic high (Figure 23). Again, it must be emphasized that the velocity inversion at the top of the upper mantle, used in the model to block long-range propagation beneath the high, is not the important feature here: many reasonable alternatives exist that would achieve this same result. However, the lowered velocities at the base of the crust are well-constrained features


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RANGE (krns) Fig. 21. Comparison of observed (water path corrected) travel

times and travel times calculated by ray tracing through the structure defined in Figures 16-18 for data north and south of OBH 1 and south of OBH 2.

and indeed are the only practical explanation that could be determined to account for the clear 7.0 km/s phase velocity observed in the range window 30-45 km south of OBH 2.

DISCUSSION AND CONCLUSIONS

The two most important general results of this work (Figure 23) may be simply stated as (1) normal oceanic crustal thicknesses underlain by 8.0 km/s upper mantle velocities exist beneath the present-day Mid-Atlantic Ridge median valley, and (2) major along-axis changes in crustal velocity structure, crustal thickness, and upper mantle velocity occur beneath a 120-km-long segment of the median valley south of the Kane Fracture Zone. These two fundamental observations have

wide-ranging implications with regard to the processes of crustal accretion and evolution at a slow spreading ridge.

The presence beneath the ridge axis of normal upper mantle material, a well-developed Moho transition zone, and a crustal velocity structure similar to that of normal young crust is compelling evidence for the cyclic nature of the crustal accretion process. Our preferred interpretation is that the along- axis topographic high upon which OBH 1 is located marks the middle of an independent spreading center cell or segment that has reached a "mature" state, that is, it has cooled suf-

ficiently since its last major episode of volcanic injection that it has taken on most of the primary characteristics of ocean crust. We use the term "spreading center cell" to define the primary unit of the segmented North Atlantic accreting boundary as described by Schouten and Klitgord [1982] and Schouten et al. !-1985a]. If the period between phases of volcanic injection into this cell is 10,000-50,000 years [Ballard and van Andel, 1977; Schouten et al., 1985a], then our results require a very efficient cooling mechanism that can operate throughout the crustal column on this time scale. Models of hydrothermal circulation systems and geochemical analyses support the presence of hydrothermal circulation deep into, and perhaps throughout the crust, and prove the capability of this mechanism for rapid cooling [e.g., Gregory and Taylor, 1981; Fehn et al., 1983; Mottl, 1983]. The primary seismic characteristics are frozen into the ocean crust in the median

valley within a few thousand years of its initial emplacement. Several tens of kilometers of the ridge axis is underlain by

this "normal" crust; the similarity between the OBH 1 and OBH 2 travel time and amplitude data and the presence of clear, second-order, intracrustal multiples propagating along the complete length of the ridge segment to OBH 2 (Figure 4c) means a considerable degree of lateral homogeneity must exist. On the scale of our seismic measurements, the processes responsible for creation of this velocity structure are apparently systematic and orderly.

A fundamental difference between the structure of this

"normal" crust underlying the ridge segment south of the 23ø15'N boundary and the structure of several million-year- old crust on the adjacent ridge flanks [Detrick and Purdy, 1980; Cormier et al., 1984] is the presence of unusually low velocities, particularly in the uppermost part of layer 3 that are in part responsible for the median valley section having a crustal delay time (1.5 s) that is several tenths of a second greater than normal. If the assumption is made that this struc-

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RANGE (rns) Comparison of observed (water path corrected) travel

times and travel times calculated by ray tracing through the structure defined in Figures 16-18 for data south of OBH 6 and south of OBH 3.


2

•0

DISTANCE (t•m•)

'1'10 '100 90 80 70 60 50 40 30 20 10 i i i i i I i i i i I I i i i I i i i i i

-SOUTH VERTICAL EXAGGERATION '-' 5:t NORTH-

•/,....• SPREADING CENTRE CELL BOUNDARY ZONES • L RIDGE • • RIDGE • RIDGE - - SEGMENT • _ SEGMENT • • SEGMENT

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Fig. 23. Cartoon illustrating principle features of along-axis changes in structure beneath the MAR median valley south of the Kane Fracture Zone.

ture will evolve into the more frequently observed velocity- depth section that has a 3- to 4-km thickness of 6.8-7.0 km/s material and a crustal delay time of ,-• 1.0 s, then a simple model of rapid crustal evolution is suggested. Three velocity depth functions are shown overlain in Figure 24 representing (1) the structure beneath the along axis topographic high (zone 4 in the ray trace model), (2) the structure of the remainder of this ridge segment (zones 3 and 5), (3) the structure on the flanks of the ridge as determined by Cormier et al. [1984]. If the important assumption is made that the major differences between these velocity-depth functions are caused only by time-dependent processes, then we associate the times t•, t2, and t3, respectively, with each of these structures (see Figure 24). The major change can be seen to be the increase in the velocity of layer 3: at time t• it has velocities less than 6.5 km/s and no Moho boundary exists. By time t 2 a well- developed Moho has been formed, and the velocity at the base of layer 3 has increased to 7 km/s. By time t3, velocities throughout layer 3 have risen to 6.8-7 km/s, and the classical oceanic velocity structure exists.

An important constraint on the process that might be responsible for this evolution is the fact that this increase in layer 3 velocities appears to begin at the base of the crustal column and to propagate upward. This argues against temperature effects being an important factor as seawater must be the primary cooling agent, and thus following a volcanic injection phase, the major temperature changes must begin at the seafloor and propagate downward through the crustal column. We suggest the following senario to account for this evolution. The time t• represents the phase immediately following the freezing of the magma chamber: Layer 3 is pervasively cracked throug• out its thickness with porosities of several

percent, sufficient to lower the velocity by 0.5 km/s or more, and active hydrothermal circulation is occurring, rapidly cooling the upper mantle which is the primary heat source now that the latent heat from the solidification of the magma chamber has been dispersed. At time t2, alteration of the low- ermost gabbros has sealed the cracks at the base of layer 3 sufficient to increase velocities to values of 6.8-7.0 km/s, and permeabilities have decreased sufficiently to restrict the active circulation to the uppermost part of layer 3 and layer 2. By time t 3 this crack-sealing process has progressed upward through layer 3, leaving behind it the 3- to 4-km-thick homogeneous 6.8-7.0 km/s layer of low-velocity gradient that is one of the most characteristic features of oceanic crust. Only the extrusives retain permeability values sufficient to support ongoing circulation, and this may continue for tens of millions of years slowly increasing the uppermost layer 2 velocities from values of 3 km/s or less up to 4.5-5.0 km/s [Houtz and Ewing, 1976]. One difficulty with this model is that it requires, albeit for a short period of time, open cracks to exist at depths greater than 2-3 km into the crust. The degree of cracking (porosity) required to decrease the lower crustal velocities by 0.5-0.7 km/s beneath the along-axis high (see depth range 5-6 km, zone 4 in Figure 17) is strongly dependent upon the ge- ometry and aspect ratio of these cracks (see, for example, Spu- dich and Orcutt [1980]). In the case of infinitely flat disk- shaped pores, or other configurations close to the Hashin- Shtrickman lower bound, the relationship between compressional wave velocity and porosity is highly nonlinear [Spudich and Orcutt, 1980, Figure 13]: an initial porosity of less than 0.5% can lower the velocity by as much as 0.5 km/s. Perhaps the lowered velocities beneath the along-axis high are simply caused by very low crack densities. We cannot prove the me-

3760 PURD¾ AND DETRICK' MID-ATLANTIC RIDGE SEISMIC STRUCTURE

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1.0

2.0

3.0

4.0

5.0

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8.0 ß i i

2.0 3.0 9.0

• , • • I , I , II

4.0 5.0 6.0 7.0 8.0

VELOCITY (krns/sec)

Fig. 24. A comparison of three velocity depth functions: the dashed line (marked t3) is the solution from Detrick and Purdy [1980] for the shots north of OBH 1 (in the 1980 experiment) located along a flow line on 7-m.y.-old crust (see location in Figure 1). The solid line (marked t2) is the best solution for OBH 1 from this experiment exactly as illustrated in Figure 8b. The chain dashed line (marked t•) is the structure of the anomalous zone beneath the along-axis topographic high. See text for discussion of possible relationships between these structures.

chanical feasibility of that explanation here, but the discovery of open cracks existing under hydrostatic pressures of 3 kbar at the Kola deep drill hole [Kozlovsky, 1982], suggests that it is not impossible.

Further speculation is required in order to associate specific values with the times t•, t2, and t 3. If time zero is taken to be the moment of initial formation of the magma chamber, then t• may be a few hundred years [-e.g., Strens and Cann, 1982'], t2 may be a few thousand years (i.e., less than the injection cycle time), and t 3 could be as little as a few tens of thousands of years. Such detailed time sequences can only be resolved by carrying out more detailed along-axis measurements, as there is little hope of detecting such rapid evolutionary patterns orthogonal to rise axes (on slow spreading ridges where the cyclicity of accretion process is presumably dominant) because at 1 cm/yr half rate, only 100 m (i.e., fraction of a seismic wavelength) of new crust is generated in 10,000 years.

Significant differences exist between the rift valley structure determined in this study at 23øN and Fowler's [1976, 1978] Fowler and Keen's [1979] results at 37øN and 45øN and Bunch and Kennett's [-1980] interpretation of the Reykjanes Ridge structure at 60øN. The crust in the median valley at 23øN is significantly thicker, and upper mantle velocities are substan- tially higher than inferred in these other studies. We interpret these differences as primarily reflecting different phases in the accretionary cycle described above. The structures determined at 37øN and 60øN, with their anomalously low upper mantle velocities and poorly developed layer 3, may correspond to our t• structure, a time only a few hundred years since the last

major volcanic event. At 23øN, on the other hand, a few thousand years may have elapsed since the beginning of the last volcanic cycle, and the structure is much closer to that of normal oceanic crust. Thus it seems meaningless to speak of "the structure of the Mid-Atlantic Ridge" unless some refer- ence is made to the evolutionary stage of the segment being studied. This is an unusual seismological problem where velocity structure is dependent not only on the three primary spatial coordinates but also upon time on a scale of a few thousand years or less.

Although the ridge segment south of the 23ø15'N boundary is predominantly homogeneous, the data define two structural anomalies both of which are restricted to a zone of a few

kilometers in width centered beneath the along-axis topographic high. The most striking of these is, of course, the zone of lowered velocities at the base of the crust. No evidence

exists for the presence of velocities as low as those previously associated with mid-ocean ridge crustal magma chambers [e.g., Orcutt et al., 1976]. It is difficult to rule out this possi- bility completely, but it seems unlikely that a zone of several kilometers in extent with velocities as low as ,-, 5 km/s could exist without the creation of a shadow zone of sufficient extent

to be observable in our data.

We interpret this zone of lowered crustal velocities to be supporting evidence for a three-dimensional model of magmatic accretion in seafloor spreading centers [Francheteau and Ballard, 1983' Whitehead et al., 1984]. We infer that the most recent crustal magma formed within this spreading center segment existed beneath this along-axis topographic high (the heating associated with this injection probably being the cause of the uplift that created the high [Francheteau and Ballard, 1983]). The zone of lowered velocity is the remnant signature of the frozen magma chamber and its feeder conduits' the zone is pervaded by cooling cracks that increase the porosity and decrease the velocity. As previously mentioned, we interpret this section of the ridge segment to be the "youngest." Perhaps a second manifestation of the youthful state of the crust around the topographic high is the unusual shallow velocity structure characterized by velocity gradients of only 0.5 s-• or less in the uppermost 500 m. The abrupt velocity discontinuity at the base of this layer lying at a depth below the seafloor of 480 m is tantalizingly coincident with the 2-3 km/s velocity increase between 450 and 550 m below the seafloor measured by the compressional wave velocity log carried out at International Program of Ocean Drilling site 395 located on 10-m.y.-old crust approximately 100 km west of OBH 1 [Hyndman and Salisbury, 1984].

The abrupt structural change at the 23ø15'N boundary is more difficult to understand. Three reasonable explanations exist for this major structural boundary and the presence of a 30-km-long ridge segment underlain by thinned and disorderly crust. Perhaps the most obvious is that because it is adjacent to the Kane Fracture Zone, this segment must be functioning with a considerably modified thermal environment because of the proximity of cold 10-m.y.-old lithosphere to the north. This must modify the nature of the accretion process in some way that it is difficult to quantify. The shortcomings of this model are its inability to explain the abruptness of the structural discontinuity at the 23ø15'N boundary (any thermal effect would presumably be gradual), and second, it would predict that this structural change would be frozen in at the time of injection and thus should be observable on the older crust. That this is not the case is well-established by three refraction lines that cross the flow line extension of the

23ø15'N boundary east of the MAR over 7-, 10-, and 15-m.y.-


old crust, respectively [Derrick and Purdy, 1980; Corntier et al., 1984].

Second, it is possible to think of this difference in structure as being due simply to this ridge segment being in a completely different cycle of the injection process: perhaps it has been starved of magma for an extended period of time and has undergone extensive stretching which has thinned the crust by 2-3 km and destroyed the orderly layering left behind by the magma chamber, resulting in the disorderly amplitude patterns seen on OBH 6. No independent evidence exists to support this model that naturally predicts the generation of inhomogeneous crustal ribbons with 2-3 km of Moho topography. There is little or no evidence to support the existence of such variability in mature oceanic crust [e.g., Purdy, 1983; White and Purdy, 1983; Diebold and Mutter, 1984] so although this "different time cycle" concept may be a contributory factor to the structural difference, it seems unlikely it is the only explanation. The third and final explanation may lie in some local- ized tectonic event. Evidence for this exists in both the nona-

lignment of the central magnetic anomaly with the median valley in this section immediately south of the Kane [Purdy et al., 1978] and the truncation of the crestal mountains north of latitudes 23ø10'N (Figure 2). In this model the thin crust and disorderly structure would be due to the accretion center just beginning to create new crust in the cold 100,000- to 200,000- year-old lithosphere. Such a ridge jump is in the same sense as the asymmetric spreading identified by Schouten et al. [1979] to have been taking place for the past several million years along this ridge segment south of the Kane Fracture Zone: 11 mm/yr to the east and 17 mm/yr to the west. Although this last explanation seems perhaps the most plausible, it is clear we have insufficient data to reach a firm conclusion. Eluci-

dation of the tectonic evolution of this ridge segment immediately south of the Kane Fracture Zone and the understanding of the formation of complex crust beneath it will come from a more detailed analysis of the morphology available from multi-narrow-beam echo-sounding data I-Detrick et al., 1984] and by using the recently collected magnetics data to more precisely define the characteristics of the central magnetic anomaly [Derrick et al., 1984].

SUMMARY OF CONCLUSIONS

1. With the exception of slightly low layer 3 velocities, the crust beneath several tens of kilometers of the MAR median

valley at 23øN has all the characteristics of simple mature ocean crust including a well-defined Moho transition zone, ,-, 8 km/s upper mantle velocities, and a total crustal thickness of 6-7 km.

2. No evidence is found for the presence of a steady state axial magma chamber in the crust or upper mantle.

3. A considerable degree of lateral homogeneity exists in the crust and upper mantle beneath the median valley over along-axis distances of several tens of kilometers.

4. These three results imply the presence of an efficient cooling mechanism operating throughout the crustal section capable of freezing-in the basic seismic characteristics of the ocean crust in a time period less than that separating the major volcanic injection events (10,000-50,000 years). Hy- drothermal circulation penetrating to the base of the crust is one plausible mechanism by which this cooling may be achieved.

5. The lower layer 3 velocities mentioned in our first conclusion may be caused by the cooling cracks that provide the permeabilities necessary to allow hydrothermal circulation to penetrate deep within the crust. Alteration products rapidly

(perhaps 10,000-100,000 years) seal these cracks, restricting circulation to the extrusive layer, reducing the layer 3 porosities, and producing typical 6.8-7.0 km/s velocities.

6. Along the 100-km-long segment of ridge studied here, two major along-axis structural changes occur. It is clear the MAR is far from being a two-dimensional structure, and it is meaningless to define the "structure of the MAR" without simultaneously defining the state and tectonic environment of the ridge segment being studied.

7. The most significant along-axis structural feature is the zone, extending 10-15 km along-axis, of lowered velocities in the lower crust centered beneath the major along-axis topographic high. This is interpreted to be the remnants of the most recent phase of injection that has temporarily left behind it a region of elevated temperatures and pervasive cracking and thus reduced velocities.

8. The second along-axis structural change occurs abruptly and catastrophically at latitude 23ø15'N and is coincident with both a major change in rift valley and crestal mountain morphology and the apparent noncoincidence of the median valley and central magnetic anomaly. North of this boundary the •, 30-km-long ridge segment that abuts the Kane Fracture Zone has a crustal thickness of 4-5 km and no

distinctive layering to produce the characteristic amplitude patterns that typify mature oceanic crust (and that are so clearly identifiable over the southernmost ridge segment). In- terpretation of the cause of this structural change is uncertain but may be related to local tectonic events, perhaps a recent 10-20 km ridge jump to the east.

Acknowledgments. We thank the officers, crew, and scientific com- plement of R/V Knorr during cruise KN92-1 when these data were collected with such care and efficiency. The excellent performance of the ocean bottom hydrophone system was due to the skill of Donald Koelsch and Carleton Grant. The large quantity of tedious data reduction required by this research was expertly carried out by Marie- Helene Cormier and Dickson Allison using programs largely written by Leon Gove. This research was sponsored by the National Science Foundation under grant OCE8025206. Woods Hole Oceanographic Institution contribution 5918.

REFERENCES

Aumento, F., B. D. Loncarevic, and D. I. Ross, Hudson geotraverse, geology of the Mid-Atlantic Ridge at 45øN, Philos. Trans. R. Soc. London, Set. A, 268, 623-650, 1970.

Ballard, R. D., and Tj. van Andel, Morphology and tectonics of the Inner Rift Valley at latitude 36ø50'N on the Mid-Atlantic Ridge, Geol. Soc. Am. Bull., 88, 507-530, 1977.

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R. S. Detrick, Graduate School of Oceanography, University of Rhode Island, Narragansett, RI 02882.

G. M. Purdy, Woods Hole Oceanographic Institution, Woods Hole, MA 02543.

(Received February 28, 1985; revised October 31, 1985;

accepted November 1, 1985.)

crustal structure of the mid-atlantic ridge at 23°n from seismic refraction studies

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