Crustal transect from the North Atlantic Knipovich Ridge to the Svalbard Margin west of Hornsund

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Tectonophysics 378 (2004) 17–41

Crustal transect from the North Atlantic Knipovich Ridge to the

Svalbard Margin west of Hornsund

Frode Ljonesa,*, Asako Kuwanob, Rolf Mjeldea, Asbjørn Breivika, Hideki Shimamurac,Yoshio Muraic, Yuichi Nishimurac

aDepartment of Earth Science, Institute of Solid Earth Physics, University of Bergen, Allegt. 41, N-5007 Bergen, NorwaybResearch Center for Prediction of Earth and Volcanic Eruptions, Tohoku University, Aramaki-aza Aoba, Aobaku,

Sendai, Miyagi 980-8578, Japanc Institute for Seismology and Volcanology, Hokkaido University, Sapporo 060, Japan

Received 12 September 2002; accepted 9 October 2003

Abstract

The crustal structure along a 312 km transect, stretching from the axial mountains of the North Atlantic Knipovich Ridge to

the continental shelf of Svalbard, has been obtained using seismic reflection data and wide angle OBS data. The resulting

seismic Vp and Vs models are further constrained by a 2-D-gravity model. The principal objective of this study is to describe and

resolve the physical and compositional properties of the crust in order to understand the processes and creation of oceanic crust

in this extremely slow-spreading counterpart of the North Atlantic Ridge Systems. Vp is estimated to be 3.50–6.05 km/s for the

upper oceanic crust (oceanic layer 2), with a marked increase away from the ridge. The measured Vp of 6.55–6.95 km/s for

oceanic layer 3A and 7.10–7.25 km/s for layer 3B, both with a Vp/Vs ratio of 1.81, except for slightly higher values at the ridge

axis, does not allow a clear distinction between gabbro and mantle-derived peridotite (10–40% serpentized). The thickness of

the oceanic crust varies a lot along the transect from the minimum of 5.6 km to a maximum of 8.1 km. The mean thickness of

6.7 km for the oceanic crust is well above the average thickness for slow-spreading ridges ( < 10 mm/year half-spreading rate).

The areas of increased thickness could be explained by large magma production-rates found in the zones of axial highs at the

ridge axis, which also have generated the off-axial highs adjacent the ridge. We suggest that these axial and off-axial highs

along the ridge control the lithological composition of the oceanic crust. This approach suggests normal gabbroic oceanic crust

to be found in the areas bound by the active magma segments (the axial and off-axial highs) and mantle-derived peridotite

outside these zone.

D 2003 Elsevier B.V. All rights reserved.

Keywords: Knipovich Ridge; Crustal structure; Oceanic crust; Vp/Vs ratio; Lithology; Slow-spreading ridge

1. Introduction

Numerous studies west of Svalbard over the last

three decades (e.g. Sundvor and Eldholm, 1979;

0040-1951/$ - see front matter D 2003 Elsevier B.V. All rights reserved.

doi:10.1016/j.tecto.2003.10.003

* Corresponding author.

E-mail address: [email protected] (F. Ljones).

Eldholm et al., 1984; Myhre, 1984; Eiken and Aus-

tegard, 1987; Myhre and Eldholm, 1987; Crane et al.,

1988; Austegard and Sundvor, 1991; Faleide et al.,

1991, 1996; Hjelstuen et al., 1996) have provided

substantial knowledge about the tectonic evolution

and physical properties of the crust in this area.

However, the majority of these investigations were

F. Ljones et al. / Tectonophysics 378 (2004) 17–4118

generally restricted to the sedimentary and deep

crustal structures in the vicinity of the western Sval-

bard Margin. The structure of the oceanic crust

generated by the slow-spreading Knipovich Ridge

still remains a relatively uninvestigated area compared

to the other North Atlantic spreading ridges further

south. The complexity of the Knipovich Ridge with

its poorly developed magnetic anomaly pattern,

oblique slow-spreading and segmentation makes this

end-member of Spreading Ridge Systems an impor-

tant and interesting ridge to investigate.

During the summer of 1998, 10 ocean bottom

seismometer (OBS) profiles with a total length of

Fig. 1. Left: main structural elements in the North Atlantic–Arctic Region a

Ridge. W.J.M.F.Z. =West-Jan Mayen Fracture Zone; M.R. =Molly Ridge

Right: zoomed section showing the bathymetry along the northern Knipov

highs are concentrated. The highs disappear in the bathymetry measuremen

direction, the distribution of off-axial highs can be traced down to the O

Neumann and Schilling (1984) where fresh basalts were found. Well DS

Depth in meters. Source: IBCAO ship track bathymetry grid for the arcti

2561 km were acquired along the western Svalbard

Margin and the northeastern Barents Sea Margin.

The experiment was performed by the Institute of

Solid Earth Physics, University of Bergen in coop-

eration with the Norwegian Petroleum Directorate

(NPD) and the Institute for Seismology and Volca-

nology, Hokkaido University, Japan. The objective of

the OBS-98 Project was to map the regional crustal

and upper mantle structures in these areas by the use

of OBS data. Several studies have shown that OBS

data are very well suited for mapping deep crustal

and upper mantle structures (e.g. Digranes et al.,

1996, 1998; Holbrook et al., 1994; Mjelde et al.,

nd the study area (within box) at the ultra-slow spreading Knipovich

; S.F.Z. = Spitsbergen Fracture Zone; M.F.Z. =Molly Fracture Zone.

ich Ridge. The outlined regions represent zones where the off-axial

ts as they are covered with sediment. Assuming constant spreading-

BS profile. White triangles represent dredge samples performed by

DP 344 indicated basalts (perhaps olivine basalts) (Talwani, 1978).

c regions.

F. Ljones et al. / Tectonophysics 378 (2004) 17–41 19

1992, 1997, 1998; Mjelde and Sellevoll, 1996;

Raum, 2000).

This paper focuses on the oceanic crust in a slow-

spreading ridge environment along a 312-km-long

transect stretching from the axial mountains of the

Knipovich Ridge, across the COT, to the continental

shelf of western Svalbard near Hornsund (Fig. 1). The

velocity models, in terms of Vp and Vs, have been

obtained by 2-D kinematic (travel time) ray-tracing

modelling and inversion of the interpreted OBS data.

The velocity distribution, including a 2-D gravity

model, is used to construct a geological model of

the structures in the continental and oceanic crust. In

addition, we analyze where P-to-S conversions occur

because this is vital for the estimation of Vp/Vs ratios

in the S-wave modelling.

2. Geological framework

The tectonic evolution of the western Svalbard

margin (north of 74jN) is largely dominated by two

a stage Cenozoic evolution. Continental break-up and

initiation of sea floor spreading between Norway and

Greenland, although not for the western Svalbard

Margin, started in the late Paleocene/early Eocene

(between anomaly 25/24) (Talwani and Eldhom,

1977). This corresponds to a geological age of 54.8

Ma according to the Gradstein and Ogg (1996) time

scale, which we use throughout the paper. During

Eocene time, between magnetic anomaly 25/24 and

13, Greenland moved in a north–northwest direction

relative to Eurasia opening the southern part of the

Norwegian-Greenland Sea. In the north, Greenland

slid along Svalbard generating the Spitsbergen Shear

Zone, a regional continental shear zone, which acted

as a plate boundary (Talwani and Eldhom, 1977). The

transpressional regime within this zone created the

Spitsbergen Orogeny (West Spitsbergen Fold and

Thrust Belt) (Harland, 1969; Eiken and Austegard,

1987). In the early Oligocene (anomaly 13, 33.7 Ma),

the relative plate motion changed to west–northwest,

causing oblique extension along the margin and

opening of the northern Greenland Sea. This was

caused by cessation of sea floor spreading in the

Labrador Sea in early Oligocene when Greenland

became part of the North American Plate (Kristof-

fersen and Talwani, 1977).

In the early Oligocene, the oblique extension

(transtension) in the Spitsbergen Shear Zone lead

to the initiation of sea floor spreading along the

Knipovich Ridge (Talwani and Eldhom, 1977). The

sea floor spreading started in the south, and gener-

ation of new oceanic crust propagated northward

with time. According to Eldholm et al. (1994)

complete continental separation along the western

Svalbard margin was not achieved before the middle

or late Miocene (about 10–15 Ma). The Hornsund

Fault Zone, which acted as a part of the Spitsbergen

Shear Zone during Eocene, became a rifted margin

in the early Oligocene.

2.1. The Knipovich Ridge

The Knipovich Ridge is the northernmost member

of the North-Atlantic spreading ridges. The ridge

starts in the south at the Mohns Ridge (f73j50VN), and approaches the continental margin of

western Svalbard northwards, where it ends in the

Molloy Fracture Zone (f 78j30VN) (Fig. 1). The

plate boundary is oblique to the spreading direction,

with obliqueness changing along strike (Eldholm et

al., 1990). The poorly developed magnetic anomaly

pattern in the oceanic crust precludes determination of

age and spreading rates from the magnetic anomaly

time scale. Estimates of spreading rate using heat

flow data by Crane et al. (1988) suggested half

spreading rates much less than 10 mm/year, with

4.3–4.9 and 1.5–3.1 mm/year for the latitudes

75jN and 78jN, respectively. However, later exami-

nations by Crane et al. (1991) suggest that the

spreading is strongly asymmetric at a rate of approx-

imately 7 mm/year to the northeast and 1 mm/year to

the southwest, indicating that the North American

Plate is moving faster than the relatively stationary

Eurasian Plate.

The ridge-valley depth is normally between 3.2

and 3.4 km, although it locally reaches depths of

more than 3.7 km (Eldholm et al., 1990; Crane et al.,

2001). The ridge is cut into segments by several

highs decreasing the rift-valley depth in these zones

by several hundred meters. Volcanic chains of off-

axial highs (Seamount Belts) are located adjacent to

these axial-highs. These volcanic active features are

well depicted in high-resolution bathymetry maps

along the spreading ridge (Fig. 1). According to

Fig. 2. Vertical high-gain component of OBS 4 and 6, 5–12 Hz band-pass filter and 2 s AGC-window applied. Reduction velocity, 8.0 km/s.



Crane et al. (2001) these features show that the

evolution of the ancient Spitsbergen Shear Zone

played a major role in the development of the

Knipovich Ridge and its segmentation. The axial-

and off-axial highs follow highly oblique strike-slip

faults, which in turn follow ‘‘zones of weakness’’

inherited from the evolution of this shear zone.

ni¼1

Ui

3. Data acquisition and processing

The acquisition of the data along the profile was

performed in August 1998 using R/V Hakon Mosby,

University of Bergen. Four Bolt 1500 C air guns with

a total volume of 77.66 l were used as a seismic

source. The shots were triggered by a Differential-

GPS navigation system with one shot every 200 m. A

single-channel streamer, with a recording length of 6 s

and sample rate of 1 ms, was used along the profile. A

LaCoste and Romberg sea gravity meter acquired the

gravity data.

Ten analogue OBS instruments were used to record

the seismic data. Two OBSs, OBS 3 and OBS 8, did

not yield useful data. The OBS instruments have three

orthogonally mounted geophones; one vertical and

two horizontal. The analogue OBSs, developed at the

Hokkaido and Tokyo universities, can record contin-

uously for 14 days with a 1–30 Hz bandwidth (� 3

dB). The data presented in this study are high-gain

versions.

The three-component OBS data were a/d converted

at the University of Hokkaido, Japan, and further

processing has been performed at the Institute of Solid

Earth Physics, University of Bergen. The processing

of the vertical components involve construction of a

band-pass filter and an AGC-window (amplitude

scaling). The OBS data are of good quality, but

contain some low-frequency noise ( < 4 Hz) short-path

ringing and water column multiples. Several band-

pass filters were tested in order to attenuate this noise

and hence to increase the signal-to-noise ratio. A

band-pass filter of 5–12 Hz was chosen together with

a 2 s AGC-window to process the data (Fig. 2). These

processing parameters were also used on the horizon-

tal components (Fig. 3). Predictive deconvolution was

also tested on the data set. A 120 ms prediction length

(gap) and varying operator length of twice the water

depth (ms) for each OBS + 120 ms were used as

processing parameters. The objective of this proce-

dure was to compress the primary signals and atten-

uate the (long-path) water column multiples in order

to help in the identification of events.

4. Description of reflection data

Constraints for the initial velocity model was

obtained from coincident reflection seismic data,

consisting of MCS line SVA-3 and the SCS line

recorded with the OBS profile (Fig. 4). SVA-3 was

acquired in 1987 by Mobil E and P Services in

cooperation with the Institute of Solid Earth Physics,

University of Bergen (Eiken and Austegard, 1987).

SVA-3 is coincident with the eastern part of the OBS

profile (145–300 km) and covers the continental shelf

and the main sedimentary basin west of Hornsund,

Svalbard. The single channel profile (0–145 km)

covers the western part of the sedimentary basin and

the Knipovich Ridge. The last 12 km in the east

(300–312 km) of the OBS profile were not mapped

by reflection data, horizons from SVA-3 were extrap-

olated into this area. The interpretation of the reflec-

tion data is based on the results presented by Fiedler

and Faleide (1996), Faleide et al. (1996) and Hjelstuen

et al. (1996).

5. Crustal modelling

5.1. P-wave modelling

The interpretation of the compiled reflection line

was depth-converted using interval velocities (Table

1) obtained from previous studies in the area. The

depth-converted version is used as a base for the 2-D

kinematic ray-tracing modelling of the OBS data. The

program package Xrayinvr, a 2-D kinematic ray-

tracing and inversion program, was used for the

velocity modelling (Zelt and Smith, 1992).

The program package includes a method for eval-

uating the goodness of fit between the observed and

calculated arrival times called v2 (chi-squared value);

v2 ¼ 1Xn T0i � Tci� �2

; ð1Þ

Fig. 3. Horizontal high-gain component of OBS 4 and 6, 5–12 Hz band-pass filter and 2 s AGC-window applied. Reduction velocity, 8.0 km/s.


Fig. 4. Interpretation of reflection seismic data. The multichannel line SVA-3 yields reliable information on deeper sedimentary layering. The

section includes Cenozoic sediments from Oligocene to present (33.6–0 Ma). Based on Hjelstuen et al. (1996) layer 2 correspond to the glacial

Sequence GIII (age 0–0.44 Ma), layer 3 to Sequence GII (age 0.44–1.0 Ma) and layer 4 to Sequence GI (age 1.0–2.3 Ma). Layers 5, 6 and 7

correspond to the pre-glacial Sequence GO (age >2.3 Ma). See Hjelstuen et al. (1996) for more detailed information.


where T0 is the observed arrival time, Tc is the

calculated arrival time, U is the estimated pick uncer-

tainty (50–100 ms) and n is the number of picks for

each phase.

The v2 method weighs the mismatch between the

observed and calculated arrival times. A value of 1 or

lower per phase indicates a good fit. However, it is

important to stress that this is not a unique indication

of the goodness for the regional model itself. Arrivals

that are weak and difficult to pick often have a larger

uncertainty (a higher chi-squared value). 3-D effects

(out of plane ray-paths), local structures unresolvable

in the model, may preclude an apparent ‘‘perfect’’ fit

(e.g. v2 < 1).The uncertainty of the interpretation (maximum

misfit between the observed and calculated travel

time) is set to 50 ms for the sedimentary and upper

crust arrivals, while deeper crustal and upper mantle

arrivals were assigned an uncertainty of 75 and 100

ms, respectively.

5.1.1. Sedimentary P-wave velocities

Estimated P-wave velocities and ray coverage for

the sedimentary arrivals are shown in Fig. 5. The

glacial sediments (0–2.3 Ma, according to Hjelstuen

et al., 1996 constrained by regional seismic strati-

graphic interpretations), including layer 2, 3 and 4,

have P-wave velocities ranging from 1.75 km/s at the

sea floor and increase up to the maximum value of

2.80 km/s in the bottom of layer 4. The pre-glacial

sediments (>2.3 Ma, according to Hjelstuen et al.,

1996) have P-wave velocities from the minimum

Table 1

Interval velocities used for the depth conversion of the reflection

seismic line based on results obtained by Myhre (1984), Austegard

and Sundvor (1991), Hjelstuen et al. (1996)

Interval Sequence Age (Ma) Velocity

(km/s)

Glacial Seawater

column

1.48

Sediment

bottom layer

GIII 0–0.44 1.9

Sediment

layer 3

GII 0.44–1.0 2.2

Sediment

layer 4

GI 1.0–2.3 2.8

Pre-glacial Sediment

layer 5

GO >2.3 3.1

Sediment

layer 6

# 3.6

Sediment

layer 7

4.1

The sequence stratigraphy is based on Hjelstuen et al. (1996).


value of 2.90 km/s at the top of layer 5 at 250 km to

4.15 km/s at the top of layer 7 at 170 km. The

modelling indicates a total sediment thickness of

about 5 km, which is in good accordance with

previous studies in the area by Eiken and Austegard

(1987), Austegard and Sundvor (1991). As depicted in

Fig. 5, the ray coverage is sparse in certain areas. No

sedimentary arrivals are observed at the axial moun-

tains in OBS 9 and 10, due to low sediment thickness.

Low ray coverage from 210–240 km is caused by the

lack of data from OBS 3. By using results from the

areas of good ray coverage, a complete velocity

distribution of the sediments can be established.

5.1.2. Crystalline P-wave velocities

Examples of the interpreted OBS data (vertical

component OBS 6) and the P-wave velocity distribu-

tion for the crystalline crust are shown in Figs. 6 and

7, respectively.

Oceanic crystalline crust is found from start of

profile to about 225 km. The course shot-interval

(200 m) makes it difficult to resolve the thin and often

irregular extrusive oceanic layer 2A, consisting of lava

flows and pillows, from the intrusive layer 2B. The

uppermost part of the oceanic crystalline crust is thus

modelled using one layer, here named oceanic layer 2.

At the ridge axis the velocity in oceanic layer 2

range from 3.50 to 4.85 km/s. There is a strong

lateral increase in P-wave velocities away from the

ridge towards the NE, whereas the Vp,top (P-wave

velocity in top of layer) towards the start of the

profile in the SW remains low (3.55 km/s). A

marked velocity increase from 5.0 km/s at f 100

km in the model to 6.05 km/s at f 110 km, west

of the large basement high at f 125–155 km, is

observed. East of the basement high, from 155 to

230 km, the velocity in the upper oceanic crust is

quite constant, ranging from 5.70 to 5.95 km/s

(Fig. 7).

The lower crust have been divided into two layers,

interpreted as oceanic layers 3A and 3B. These two

layers have relatively constant velocities ranging from

6.55–6.95 to 7.10–7.25 km/s, respectively. However,

a drop in the seismic velocities is observed at the

COT. As opposed to the upper oceanic crust, no

evident drop in P-wave velocities is observed under

the ridge axis for the lower oceanic crust. In the upper

mantle, the P-wave velocity drops from 8.00 to 7.60

km/s at the ridge axis (Fig. 7).

The continental crystalline crust extends from

about 245 km to the end of the profile. Horst

and graben structures within the Hornsund Fault

Zone (250–275 km) are prominent features of this

part of the model. The moderate data quality of

OBS 1 and OBS 2 is most likely caused by the

complex structural geology found in this fault zone,

making a good reconstruction of the continental

crust difficult to establish. Relatively low P-wave

velocities of 5.20–5.70 km/s are found in the

basement (upper crystalline crust). The ray coverage

in the deeper continental crust is limited, and only

the western part of the area is mapped. The P-wave

velocities indicated by the ray-tracing are deter-

mined to be to 6.10–6.65 km/s in the middle

crystalline crust and 6.60–6.80 km/s in the deepest

crystalline layer (Fig. 7).

5.1.3. Chi-squared value

An assembly of the chi-squared value for the

different phases identified in the P-wave model is

listed in Table 2. The basement phases (oceanic layer

2 and crystalline upper continental crust) yield the

highest chi-squared values. This is most likely caused

by the large topographic variation seen in the base-

ment along the profile, making a good fit difficult.

The chi-squared values are close to 1 for most

Fig. 5. The 2-D model for the sedimentary layers of the OBS profile, with estimated P-wave velocities (in km/s). The top and bottom velocities have been averaged in thin layers.

KR= axial valley of the Knipovich Ridge; HFZ=Hornsund Fault Zone.

F.Ljones

etal./Tecto

nophysics

378(2004)17–41

25

Fig. 6. Interpreted version of vertical component OBS 6 with ray paths of the calculated travel-time curves. Band-pass filter (5–12 Hz) and 2 s

AGC-window applied. Displayed with 8.0 km/s reduction velocity. D =water arrival, Puc = upper crust refraction, PUC = upper crust reflection,

Plc = lower crust refraction, PILC = intra-lower crust reflection, Pn =mantel refraction.


Fig. 7. The 2-D model for the crystalline layers of the OBS profile, with estimated P-wave velocities (in km/s). The top and bottom velocities have been averaged in thin layers.

F.Ljones

etal./Tecto

nophysics

378(2004)17–41

27

Table 2

Number of picks, picking error, rms misfit and v2-value for phases

identified along the profile in the P-wave modelling

Phase Number

of picks

Picking

error (s)

rms misfit

(s)

v2

Sedim. Refrac.

sed. layer 2

28 0.050 0.058 1.373

Refrac.

sed. layer 3

28 0.050 0.049 1.006

Refrac.

sed. layer 4

17 0.050 0.058 1.451

Refrac.

sed. layer 5

26 0.050 0.050 1.057

Refrac.

sed. layer 6

48 0.050 0.049 0.979

Refrac. 9 0.050 0.050 1.103

sed layer 7* 156

Ocean. Layer 2 180 0.050 0.072 2.070

Layer 2/layer

3A refl.

107 0.050 0.052 1.081

Intra-refl. layer

3B refl.

14 0.075 0.032 0.197

Layer 3A/layer

3B refl.

56 0.075 0.069 0.872

Layer 3A 305 0.075 0.065 0.755

Layer 3B 215 0.075 0.093 1.557

Moho 620 0.100 0.133 1.780

Moho* 254 0.100 0.081 0.651

Moho reflection 11 0.100 0.060 0.706

1762

Cont. Upper crust* 52 0.050 0.112 5.143

Upper/mid-

crust refl.

71 0.050 0.060 1.480

Middle crust 41 0.075 0.099 1.480

Middle crust* 56 0.075 0.065 0.755

Intra-mid 14 0.075 0.061 0.702

crust refl. 234

The table is divided into sedimentary, oceanic and continental

sections. Symbol (*), refracted wave traveling as a head wave. The

water phases are not listed in the table, because the P-wave water

velocity is set to a constant value of 1.48 km/s. The water-arrival is

used as a guidance in order to test the right position between the

observed and calculated position of the OBS instruments, which for

all the OBSs are good.


phases, indicating that the velocity-distribution and

layer thickness inferred from the P-wave modelling

give a satisfactory representation of the subsurface

velocity.

5.2. Gravity modelling

The starting geometry of the gravity model was

determined by the P-wave model, and averaged P-

wave velocities for each layer are converted into

initial densities using the velocity–density relation-

ship of Ludwig et al. (1970). The gravity modelling

was performed using the Interactive 2.5D Gravity

and Magnetic Modelling Program GRAVMAG (Ped-

ley, 1993). The purpose of the gravity model is both

to test the validity of the P-wave model and to

determine the depth and density of the layers not

mapped by the ray-tracing method, for example the

deeper continental crust in the eastern part of the

model (see Fig. 7).

The final gravity model and the relationship

between the Nafe–Drake density-curve (Ludwig et

al., 1970) and estimated densities for each polygon

are shown in Fig. 8. No major changes in the initial

crustal densities or layer boundaries were needed in

order to fit the observed field. However, the large

lateral density-contrast in the upper crystalline layer

from 2.75 to 2.4 g/cm3 (250–312 km in Fig. 8) is

probably caused by a gradual decrease in the density

towards the northeast which is difficult to reproduce

in the modelling software package which does not

allow continuous density changes. The amplitude of

the observed gravity at the axial valley is well

reproduced in the model, as well as that of the horst

and graben structures at the Hornsund Fault Zone

(255–275 km). Increased sediment thickness along

the basin west of the continental margin is reflected

by the decreasing residual gravity field, and con-

strains the sediment thickness inferred from the P-

wave model. Lateral increase of the mantle density

towards the northeast, away from the ridge, is

interpreted to reflect the thermal structure in the

lithosphere, where lithospheric mantle density

increases due to cooling away from the ridge (Brei-

vik et al., 1999).

In general there seems to be a good fit between

the observed and calculated gravity field, indicating

a well-constrained crustal thickness and density dis-

tribution. Local small-scaled deviations are observed.

These misfits are most likely produced by 3-D

structures not resolved in the 2-D model, and their

amplitude remains within the uncertainty of this

method (F 10 mgal). The most prominent misfit

(20 mgal) is caused by the basement high at 140

km. The misfit suggests that the basement-high,

corresponding to the seamount chains described by

Crane et al. (2001), is a local high without large

Fig. 8. Left: 2-D gravity model for the profile based on the geometry inferred from the P-wave model. Right: the relation between the densities

used in the modelling and the Nafe–Drake curve (Ludwig et al., 1970). Densities in g/cm3. KR=Knipovich Ridge, HFZ=Hornsund Fault

Zone.


lateral continuity and/or is out of the plane of the

model.

5.3. S-wave modelling

The geometries obtained from the P-wave velocity

model are used as a starting model in the S-wave

modelling. The OBS data from the horizontal com-

ponents are modeled using the 2-D ray tracing

program to match the onset of the shear-wave arrivals

in order to obtain shear-wave velocities and conver-

sion levels. The maximum uncertainty is set to 200

ms for the shear wave-arrivals. This relatively large

uncertainty is based on two factors; firstly, difficulties

in picking first arrivals due to lower signal-to-noise

ratio for the horizontal component compared to

vertical component, and secondly the software pro-

gram makes it difficult to vary the Poisson’s ratio

values within each layer in a densely parameterized

model. Each layer therefore has a constant Poisson’s

value.

PPS and PSS are the two modes of arrivals that

are generally observed in the horizontal component

data. The PPS arrivals propagate with an apparent

P-wave velocity, and hence represent waves that

have been P-to-S converted on the way up to the

OBS. The PSS arrivals propagate with an apparent

S-wave velocity, indicating that they have been P-

to-S converted on the way down (Mjelde et al.,

2002a). The detection of the shear-wave arrivals is

based on comparing the horizontal component with

the vertical component in order to distinguish PPS

arrivals from P-wave arrivals and sea floor multi-

ples. The PSS arrivals are characterized by low

apparent velocities (high dip) in the recorded data.

All shear-wave arrivals have been modelled using a


single mode conversion along the ray-path, which is

the most likely mode of propagation (Digranes et

al., 1998). The modelling procedure is based on the

assumption that an interface for P-waves corre-

sponds to an interface for S-waves. This procedure

is needed since the quality of the P-waves are

generally higher than that of S-waves (Mjelde et

al., 2002b). An example of the interpreted data

(OBS 6 horizontal component) and the final S-wave

model, with average Vp/Vs are depicted in Figs. 9

and 10, respectively. The amount of PSS arrivals is

limited for OBS 1, 2 and 10 and no PSS arrivals

observed for OBS 9.

5.3.1. Sedimentary Vp/Vs ratios

High Vp/Vs ratios (from 2.00 to 7.14) in layer 2 to

layer 4 indicate poorly consolidated sediments just

below the sea bed, with Vp/Vs decreasing rapidly as

the sediments become more consolidated with depth.

In layer 5 to layer 7 the Vp/Vs ratio range from 1.87 to

1.78 (Fig. 10).

5.3.2. Crystalline crust Vp/Vs ratios

High Vp/Vs ratios (2.03 and 2.14) in upper

crystalline crust at the ridge axis reflect the high

porosity and fracture density in young oceanic crust.

From about 90 km along the profile to 240 km a

constant Vp/Vs ratio of 1.81 is found to fit the data

for the upper oceanic crust (Fig. 10). For oceanic

layer 3A, a Vp/Vs ratio of 2.03 west of the ridge

axis, 1.87 east of the ridge axis, and decreasing to

1.81 for the rest of the profile fit the data. No ray

coverage for the western part of oceanic layer 3B

and upper mantle leaves the S-wave velocity in this

area (0–50 km) unresolved. From 50 km and

northeastwards, oceanic layer 3B and the upper

mantle show similar Vp/Vs ratios as oceanic layer

3A (Fig. 10).

No S-wave coverage exists over the crystalline

crust at the COT (225–245 km). In the continental

crust (245–312 km), a Vp/Vs ratio of 1.71 for the

upper and middle continental crystalline layers yields

Fig. 9. Interpreted version of horizontal component OBS 6 with ray paths f

S-wave on the way down. PPS arrivals are similar to the ray paths shown in

way up. Solid lines represent P-waves, dotted lines S-waves. The large

indicate the presence of a conversion interface, which is not resolved in the

PSSUC = upper crust reflection, PSSlc = lower crust refraction, PSSILC = int

the best fit to the data. Only P-wave energy limited to

the western part (about 245–260 km) covers the

deepest continental crust, hence no Vp/Vs ratio is

found for this layer (Fig. 10).

5.3.3. Chi squared-values: S-wave modelling

An assembly of the chi-squared values for the

different phases identified in the S-wave modelling is

listed in Table 3. For most phases, the v2-value are

lower than 1. However, a relatively large misfit is

seen for the PPS arrivals in OBS 9 and 10 at the

ridge axis, indicating a conversion boundary in the

upper crust not resolved in the P-wave modelling.

This is interpreted as indirect evidence of a thin layer

of high porosity and low velocity at the top of

oceanic layer 2A.

5.3.4. P–S conversion

There seems to be a general agreement that the

majority of P-to-S conversions in large scale sur-

veys take place at the sediments/basement bound-

ary, due to the large acoustic impedance-contrast

between these interfaces (Mjelde et al., 1992).

Conversion interfaces for both the PPS (converted

to S-waves on the way up) and PSS arrivals

(converted to S-waves on the way down) are listed

in Table 3. For the PPS arrivals about 40% of the

phases are converted at internal sedimentary bound-

aries, whereas 50% are converted at the basement/

sedimentary boundary. The remaining 10% are

converted at the oceanic layer 2/layer 3A boundary

within the oceanic crystalline crust. Whereas 60%

of the PSS phases are converted within the sedi-

ments and 40% at the basement/sediment boundary.

Only for three PSS phases (OBS 7, (1) in Table 3)

the conversions occurred at the sea floor. The high

Vp/Vs ratios and low P-wave velocities observed at

the sea floor are probably the main reason why so

few P–S conversions occurred there. In areas of

thick sediments (140–250 km in Fig. 10) P-to-S

conversions within the sediment sequences are

favored, whereas P-to-S conversions at the base-

or the calculated travel-time curves of the PSS arrivals, converted to

Fig. 6 only delayed in time due to P–S conversion occurring on the

misfit seen for the upper crust reflection (PPSLC) in OBS 6 might

P-wave modelling. D =water arrival, PSSuc = upper crust refraction,

ra-lower crust reflection, PSSn =mantel refraction.


Fig. 10. 2-D regional model along the OBS profile, with estimated P-wave velocities displayed by Vp/Vs ratios and interpolated P-wave velocities (in km/s). Maximum oceanic crustal

thickness of 8.1 km is found at the basement high at 140 km and minimum thickness of 5.6 km at 180 km. KR=Knipovich Ridge, HFZ=Hornsund Fault Zone.

F.Ljones

etal./Tecto

nophysics

378(2004)17–41

32

Table 3

Number of picks, picking error, rms misfit, v2-values and conversion interface (‘‘P–S inter-f.’’) for phases identified along the profile in the S-

wave modelling

Mode Phase No. of picks rms misfit (s) v2 P–S inter-f

OBS 1 PPS Upper crust* 34 0.091 0.212 Sed. 3/4

Upper crust* 28 0.167 0.725 Sed. 2/3

Upper/mid crust refl. 15 0.102 0.280 Sed. 2/3

Middle crust 20 0.297 2.322 Sed. 2/3

Moho 28 0.217 1.226 Sed./basem.

Intra-mid crust refl. 8 0.310 2.738 Basem./mid c.

PSS Upper crust* 14 0.231 1.441 Sed./basem.

147

OBS 2 PPS Upper crust* 22 0.088 0.204 Sed. 3/4

Middle crust* 37 0.507 3.701 Sed./basem.

Moho* 24 0.386 6.698 Sed. 3/4

PSS Upper crust* 8 0.288 2.370 Sed. 5/6

Upper crust* 14 0.109 0.320 Sed./basem.



Middle crust 9 0.269 2.028 Sed./basem

125

OBS 4 PPS Layer 2 6 0.113 0.380 Sed. 2/3

Layer 3A 28 0.101 0.264 Sed. 2/3

Layer 3A 38 0.247 1.571 Sed./basem.

Moho* 16 0.111 0.331 Sed. 2/3

Layer 2/3A refl. 9 0.080 0.158 Sed. 2/3

Layer 3A/3B refl. 22 0.078 0.181 Sed. 2/3

PSS Layer 3B* 8 0.191 1.047 Sed. 5/6

Moho* 14 0.114 0.349 Sed./basem.

Moho* 9 0.050 0.072 Sed./basem.

Moho* 4 0.085 0.240 Sed. 6/7

Moho* 7 0.261 1.988 Sed. 6/7

Moho* 14 0.142 0.544 Sed. 5/6

Layer 2/3A refl. 6 0.068 0.137 Sed. 6/7

Layer 3A/3B refl. 24 0.103 0.279 Sed./basem.

Layer 2/3A refl. 13 0.109 0.320 Sed. 5/6

Moho refl. 13 0.142 0.546 Sed. 2/3

231

OBS 5 PPS Layer 2 12 0.156 0.665 Sed. 3/4

Layer 2 16 0.171 0.814 Sed./basem.

Layer 3A 35 0.161 1.403 Sed. 4/5

Layer 3B 11 0.095 0.248 Sed. 4/5


Layer 3B 11 0.099 0.268 Sed./basem.

Moho 6 0.142 0.603 Sed. 4/5

Moho* 19 0.083 0.201 Sed. 4/5

Layer 2/3A refl. 16 0.165 0.730 Sed. 3/4

Layer 2/3A refl. 20 0.175 0.804 Sed./bsem.

Layer 3A/3B refl. 19 0.125 0.411 Sed. 4/5

PSS Layer 2* 10 0.089 0.220 Sed. 5/6


Layer 3A 9 0.155 0.677 Sed. 4/5

Layer 3A 13 0.058 0.091 Sed. 4/5

Layer 3A* 8 0.040 0.047 Sed. 4/5

Layer 3B 7 0.140 0.574 Sed. 3/4

(continued on next page)


Table 3 (continued)

Mode Phase No. of picks rms misfit (s) v2 P–S inter-f

OBS 5 PSS Layer 3B* 18 0.109 0.312 Sed. 3/4

Layer 2/3A refl. 15 0.094 0.238 Sed. 4/5

Layer 3A/3B refl. 12 0.178 0.866 Sed./basem.

Moho refl. 13 0.030 0.024 Sed./basem.

290

OBS 6 PPS Layer 3A 30 0.093 0.222 Layer 2/3A

Layer 3B 39 0.070 0.126 Layer 2/3A

Moho 34 0.255 1.680 Layer 2/3A

Layer 2/3A refl. 20 0.455 5.442 Sed./basem.

Intra-layer 3 refl. 12 0.042 0.049 Layer 2/3A

PSS Layer 3A 21 0.165 0.717 Sed. 2/3

Layer 3B 24 0.252 1.655 Sed. 3/4

Layer 3A* 10 0.217 1.303 Sed. 4/5

Layer 3B* 18 0.281 2.086 Sed. 4/5

Moho* 89 0.185 0.864 Sed./basem.

Moho* 23 0.174 0.787 Sed. 3/4

Moho* 8 0.098 0.272 Sed. 3/4

Moho* 8 0.180 0.930 Sed. 4/5

Layer 2/3A refl. 12 0.231 1.461 Sed. 2/3

Layer 2/3A refl. 18 0.244 1.571 Sed. 3/4

363

OBS 7 (� 1) PPS Layer 2 12 0.380 3.911 Sed./basem.



PSS Moho* 4 0.143 0.679 Sed./basem.

Moho* 24 0.269 1.883 Sed./basem.


OBS 7 (1) PPS Layer 2 18 0.068 0.123 Sed./basem.

Moho* 33 0.075 0.146 Sed./basem.


Layer 3A* 11 0.218 1.312 Sed./basem.

Layer 3B* 9 0.048 0.065 Sed./basem.

Moho* 28 0.130 0.438 Sed./basem.

PSS Layer 3A* 12 0.197 1.056 Ocean-b.

Layer 3B* 14 0.301 2.446 Ocean-b.

Moho* 12 0.141 0.141 Ocean-b.

268

OBS 9 PPS Layer 3A 30 0.227 1.328 Sed./basem.


Moho 50 0.378 3.642 Sed./basem.


143

OBS 10 PPS Layer 2 26 0.174 0.790 Sed./basem.


Moho 67 0.204 1.052 Sed./basem.

PSS Layer 2 5 0.110 0.377 Sed./basem.


149

Each OBS listed separately in order to describe each phase in more detail. Symbol (*), refracted wave traveling as a head wave.


ment/sedimentary boundary are favored in areas of

thin sediments (0–150 km in Fig. 10). The results

presented here is in good agreement with the results

presented by Mjelde et al. (2002a), where the

distributions of P-to-S conversions follow a similar

pattern.


6. Modelling uncertainties

The uncertainty in the modelling is directly related

to the interpretative step of phase identification rather

than the actual travel-time modelling. The uncertainty

of the interpretation is unfortunately almost impossi-

ble to quantify, but is clearly related to the data

quality. P-wave velocities in seismic refraction studies

are often estimated to an accuracy of F 0.1 km/s

(Mjelde et al., 2002b) with an uncertainty in crustal

thickness of F 0.5 km in areas of good data quality

and ray coverage. The data quality and ray coverage

(see Figs. 2 and 7) is good for most of the P-wave

data, except for OBS 1 and 2 covering the continental

part of the profile which therefore has an uncertainty

larger than F 0.1 km/s. The quality and ray coverage

for the S-wave data is relatively good for OBS 4–7

and moderate for OBS 1, 2, 9 and 10. We estimate the

uncertainty in the S-wave velocities and Vp/Vs ratios

to be in agreement with those found for several data

sets further south, at the Mid-Norwegian margin and

in the Jan Mayen Basin (Digranes et al., 1996; Mjelde

and Sellevoll, 1996; Mjelde et al., 2002b). In these

areas, the uncertainty in the Vp/Vs ratios was estimated

to be F 0.05 in areas with good data quality and

F 0.07 in areas of moderate data quality. This implies

an uncertainty of F 0.05 for the data covering the

central part of the profile (70–245 km) and F 0.07

for the S-wave data covering the continental part of

the profile (245–312 km) mapped by OBS 1 and 2,

and for the western part (0–70 km) mapped by OBS 9

and 10.

7. Discussion

The seismic velocity models derived from the

OBS data together with the gravity model provide

important insights into the crustal structure along the

profile. In this section, we discuss these results

emphasizing on the crustal thickness and lithology.

7.1. Lithology in the sedimentary layers

The connection between Vp/Vs ratios and lithology

is well established (e.g. Hamilton, 1979; Domenico,

1984; Johnston and Christensen, 1992), although this

is not a unique diagnostic indicator. Variations in rock

parameters such as porosity, pore fluid, pore geometry

and degree of consolidation affect the Vp/Vs ratio in

sedimentary rocks and care must be taken in order to

constrain lithology (Tatham, 1982). The Vp/Vs ratios

are average values and lateral and vertical variations

probably exist within the intervals resolved in the

modelling. Based on laboratory studies on sedimen-

tary rocks Vp/Vs ratios range from 1.59 to 1.75 for

pure sandstones and 1.7 to 3.0 for shales (Domenico,

1984).

The Vp/Vs ratios (from 2.45 to 7.14) in layers 2 and

3 suggest high-porosity muddy sediments at the sea

floor changing to mudstone and perhaps shale within

deeper parts of layer 3. The Vp/Vs ratio of 2.00 in layer

4 suggests shale as the most likely composition in this

layer. The lithology in the pre-glacial sediments is

constrained by comparing the Vp/Vs ratios with sedi-

mentology studies by Hjelstuen et al. (1996) in the

Storfjorden Fan, south of Svalbard, in the vicinity of

the profile. According to Hjelstuen et al. (1996), the

GO Sequence (layers 5, 6 and 7 in Fig. 5) are

dominated by fluviatile drainage systems, with rivers

transporting sediments to the margin. Based on this, a

mixture of sand and shale is likely. The decreasing Vp/

Vs ratios (1.87, 1.84 and 1.78 for layers 5, 6 and 7,

respectively) supports a sand–shale composition with

sand content increasing with depth.

7.2. Crystalline crust

7.2.1. COT: continent ocean transition

The transition from oceanic to continental crust is

found within a 20 km wide zone at 225–245 km just

west of the Hornsund Fault Zone. This is marked by a

sudden increase in crustal thickness and a distinct

increase in the crystalline P-wave velocities from east

to west within this zone (Figs. 7 and 10). This

observation is in good accordance with Breivik et al.

(1999) and Myhre and Eldholm (1987) which suggest

separation of oceanic to continental crust within a

narrow zone of 10–30 km just west of the Hornsund

Fault Zone (north of 76jN).

7.2.2. Continental crystalline crust

As mentioned earlier (see Sections 5.1.2 and

5.3.2), the limited ray coverage precluded a detailed

analysis of the continental crust. The relatively low P-

wave velocity of 5.20–5.70 km/s and density of


2.40–2.75 g/cm 3 in the upper continental crust (245–

312 km in Figs. 7 and 8) suggest metasediment

(metamorphosed pre-Oligocene sediments) generated

by the thin skinned folding in the Spitsbergen Orog-

eny in late Paleocene–Eocene (Steel et al., 1985).

Extension tectonics starting in Oligocene (anomaly

13, 33.7 Ma), responsible for the horst and graben at

250–275 km in Fig. 10 (Eiken and Austegard, 1987)

and overall increased porosity and fracture density,

could explain the low P-wave velocities in the upper

continental crust. The P-wave velocity of 6.10–6.65

km/s and average density of 2.90 g/cm3 in the middle

continental crust (Figs. 7 and 8), with a Vp/Vs ratio of

1.71, indicating a fairly high quartz content (Chris-

tensen, 1996), is reflecting an average felsic rock

composition.

The continental crustal thickness is not constrained

in the velocity modelling due to the absent ray

coverage in the deepest crystal layer in the east

(250–312 km; Fig. 10). A mantle density of 3.345

g/cm 3 under the continental crust fits well to the

observed gravity data. This yield a crustal thickness

of f 24 km from 260 to 312 km, with a well defined

thinning from 22 to 14.5 km at 260–240 km towards

the COT (Fig. 8). This model is however not consis-

tent with the gravity model proposed by Austegard

and Sundvor (1991) which indicates a crustal thick-

ness of 33 km from 275–312 km along the profile,

decreasing rapidly to 18 km at 265 km in Fig. 8.

Although their model lacks information about the

initial velocities and layer boundaries for the middle

and lower continental crust, their crustal model cannot

be ruled out. The results presented in this study clearly

shows that the continental crust in this area needs to

be further investigated.

7.2.3. Oceanic crust

According to White et al. (1992), the mean thick-

ness for normal oceanic crust with a half spreading

rate of f 10 mm/year is 7.1F 0.8 km. Oceanic crust

generated at half-spreading rates above 10 mm/year

will have the same thickness because the total amount

of melt produced is independent of spreading rate

once the half-spreading rate is above 10 mm/year.

However, oceanic crust generated at half-spreading

rates below 10 mm/year will have a marked decrease

in the amount of melt generated due to conductive

heat loss from the mantle beneath the spreading axis,

and the crust will be thinner. According to Crane et al.

(1991), the half-spreading rates for the Knipovich

Ridge have been well below 10 mm/year since the

initiation of sea floor spreading in Oligocene to

present. Based on these assumptions we should expect

a thin crust ( < 6 km) in this area from the analysis of

White et al. (1992). The mean thickness of the oceanic

crust inferred from the P-wave modelling and further

constrained by the gravity model is 6.7 km, which is

thicker than expected from White et al. (1992). Our

observations are however consistent with the results

reported by Weigelt and Jokat (2001) from the ultra

slow-spreading Gakkel Ridge in the Arctic Ocean,

known to be the slowest of all Mid-Ocean Ridge

Systems. Their gravity models show a variation in

crustal thickness from 3 km up to extreme values of 9

km. According to Weigelt and Jokat (2001), this

cannot be explained by the theoretical models of

oceanic crustal formation proposed by Reid and

Jackson (1981), Bown and White (1994) and Su et

al. (1994). The large thickness of the oceanic crust at

the Knipovich Ridge could be related to the magmati-

cally active cells at the ridge axis (and the off-axial

highs adjacent to the ridge) described by Crane et al.

(2001), which have a high magmatic activity regard-

less of spreading rate.

7.2.4. Upper oceanic crust

The lateral increase in the P-wave velocity and

lateral drop in the Vp/Vs ratio in the upper crust (Fig.

10) is due to hydrothermal circulation closing cracks

and decreasing overall porosity as the crust mature

(Grevemeyer and Weigel, 1996). The upper crustal

Vp/Vs ratio of 1.81 east of the axis is in good

accordance with the results obtained by Klingelhoe-

fer et al. (2000) at the ultra slow-spreading Mohns

Ridge, south of the Knipovich Ridge. Here the

authors proposed Vp/Vs ratio of 1.78 and 1.81 for

oceanic layer 2A and layer 2B, respectively.

7.2.5. Lower oceanic crust

At the spreading axis a lateral change in Vp/Vs ratio

in oceanic layer 3A from 2.03 (west) to 1.87 (east),

and 1.87 for oceanic layer 3B just west of the axis fit

the data (Fig. 10). Despite the moderate data quality

and low ray coverage in this area compared to the

central part (70–245 km) the S-wave modelling

indicate that the Vp/Vs ratio is reflecting a high fracture


density in the crust and upper mantle at the ridge axis,

allowing seawater circulation in the lower crust and

serpentization of the ultramafic minerals (olivine and

pyroxene). The process is considered as a major factor

of metamorphism at slow-spreading ridges, where

extension of the lithosphere and lack of melt can lead

to emplacements of mantle rocks (peridotite) into the

gabbroic lower crust (Karson et al., 1987; Cannat,

1993; Mjelde et al., 2002b).

In the region between 80 and 225 km the constant

Vp/Vs ratio of 1.81 in the mature lower crust indicates

that the Vp/Vs ratio is determined by the mineral

composition. Results from Horen et al. (1996) suggest

that gabbro and peridotite with 10–40% serpentiza-

tion have identical Vp/Vs ratios of 1.73–2.08 for P-

wave velocities of 6.1–7.5 km/s. Spudich and Orcutt

(1980) have estimated the Vp/Vs ratio for gabbro and

upper mantle peridotites exposed in ophiolites to be

1.81–1.87 and 1.81–1.99, respectively.

Carlson and Miller (1997) have shown that partial-

ly serpentinized peridotites can be distinguished from

oceanic layer 3 gabbros based on differences in Vp/Vs

ratios. However, for P-wave velocities of about 6.8–

7.2 km/s, which falls within our velocity estimates in

oceanic layer 3B, the Vp/Vs ratio of 1.81 overlap both

trends. We therefore agree with Klingelhoefer et al.

(2000) that the Vp/Vs ratio of 1.81 in oceanic layer 3B

does not allow a clear distinction between gabbro and

10–40% serpentized peridotite or a small scale mix-

ing of these in the lower crust.

7.2.6. Alternative model

Ridge segments characterized by low spreading

rates, low degree of partial melting and low ascent

velocities in the asthenosphere are thought to be the

best candidates for emplacements of mantle periodites

at the ridge axis sea floor. An increase of these factors

will lower the probability of finding mantle or deep

crustal rock outcrops (Cannat, 1991). Outcrops of

mantle derived rocks normally occur along slow-

spreading ridge segments with well-developed deep

axial valleys (Cannat, 1993), which characterizes the

Knipovich Ridge. According to Phipps-Morgan et al.

(1994) deep axial valleys reflect hydrothermal cooling

as the controlling factor, at the expense of the magma

production. The deep axial valley corresponds to a

thick axial lithosphere, which reduces the ascent of

melt from the asthenosphere.

There is a contradiction between geological and

geophysical observations in areas where axial serpen-

tized peridotites are found. Geological observations

suggest zero to near zero magmatic crustal thickness,

whereas seismic and gravity results indicate a few

kilometers thick oceanic crust, with moderate densi-

ties and seismic velocities (Cannat, 1993). Cannat

(1993) explains this by mantle-derived peridotites

tectonically uplifted with gabbro-bodies added by

short-lived intrusions. The seismic velocities are not

necessarily higher than normal oceanic crust in these

magma-starved areas, because the regions are normal-

ly heavily tectonized, with faults and crushed rocks

allowing extensive serpentization of the mantle de-

rived rocks.

Gravity studies from the Mid-Atlantic (22–24jN)by Cannat et al. (1995) suggest a strong correlation

between thin oceanic crust and the emplacement of

mantle-derived ultramafic rocks (serpentized perido-

tite) in the sea floor along the ridge axis. Positive

residual gravity anomalies correspond to segments of

poor magmatic activity (corresponding to a thin crust

dominated by mantle-derived rock) and negative re-

sidual anomalies corresponding to magmatically ac-

tive segments. The model proposed by Cannat et al.

(1995) shown in Fig. 11 can be used as an analogue to

the segmentation pattern along the Knipovich Ridge,

where the zones of axial and off-axial highs corre-

spond to the magmatic active segments and the areas

between these zones to the magma-starved areas

where mantle-derived rocks are expected.

From Figs. 1 and 11 the magmatically active seg-

ments, corresponding to the axial and off-axial highs

along the profile, are expected to appear at the

spreading axis, near OBS 6 and just east of OBS 4.

The magma-starved areas should be found within a

narrow zone at OBS 7 and between OBS 5 and OBS

4. Compared with Fig. 10 these coincide well with the

high at OBS 6 described earlier which is particularly

pronounced. The high east of OBS 4 (200 km in Fig.

10), although close to the COT, is also clearly ob-

served. The magma-starved zone at OBS 7 corre-

sponds to the velocity-contrast in the upper crust

found in the P-wave modelling (Vp,top = 5.0 km/s at

100 km increasing to 6.05 km at 110 km in Figs. 7 and

10). An indication of a magma-starved zone is also

found in the gravity data over the same area as a

positive short wavelength residual anomaly, indicating

Fig. 11. Sketch showing the along-axis section of two idealized

magmatic segments. Magmatic crust is shown as continuous and

layered at magma-rich segments centers but becoming progressively

thinner and more discontinuous with arrays of short strike-slip and

oblique faults, toward magma-poor segment ends, where ultramafic

outcrops are common (after Cannat et al., 1995).


that this body could be a partly serpentized ultramafic

rock. No such velocity contrast or gravity anomaly is

found in the area between 155 and 200 km and east of

OBS 4 (200–225 km; Fig. 10). This is consistent with

the results presented by Cannat (1993), which indicate

that it is very difficult to distinguish normal oceanic

crust from serpentized peridotites based on pure

geophysical observations. However, fresh basalts have

only been dredged along the axial valley within the

same zones as sketched in Fig. 1 (Neumann and

Schilling, 1984). In the axial valley outside these

zones only sediments were found.

Although a maximum crustal thickness of 8.1 km

is found at the basement high at 140 km and minimum

thickness of 5.6 km at 180 km (Fig. 10), the models

presented here does not show an overall marked

thinner crust in the areas where mantle-derived rock

is expected to dominate. Furthermore, it is found to be

very difficult to resolve extreme local variations in

crustal thickness in the software package using a

three-layered oceanic crust model. Two alternative

models, with one- and two-layered oceanic crust, were

tested in order to elucidate this, and although they

could account for larger local variations in thickness

the two models yield a poorer fit to the data.

The results presented here show the complexity of

proving the existence of a mantle-derived crust. Al-

though the results inferred from the velocity and

gravity modelling are nonconclusive they do give

tentative indications of a complex oceanic crust which

could be correlated to an alternating normal and

mantle-derived oceanic crust. It is however important

to point out that the OBS line presented in this paper

intersects both the Knipovich Ridge and the estimated

spreading direction at an acute angle. In addition, the

large spacing between the OBSs give a poor lateral

resolution. These factors are forcing us to average the

distribution of the off-axial highs and magma-starved

segments. We therefore suggest OBS surveys to be

performed parallel to these segments in order to better

describe the oceanic thickness and crustal composi-

tion along the ridge. A survey focusing on dredging

peridotites is planned along the Knipovich Ridge

(R.B. Pedersen, personal communication), which

hopefully will give better constraints on the crustal

composition along the ridge.

8. Conclusions

Modelling of OBS data has provided new infor-

mation on the seismic velocities, thickness and com-

position of the crust along a 312-km transect west of

Svalbard. A geological model based on the inferred

results is shown in Fig. 12.

Our results suggest muddy sediments in the sea

floor changing to shale in deeper part of the glacial

sediments ( < 2.3 Ma). Decreasing Vp/Vs ratio in the

pre-glacial sediments (>2.3 Ma) indicate sedimentary

rocks with a sand–shale ratio increasing with depth.

The S-wave modelling has shown that most of the P to

S-conversions are occurring at the sediment/basement

boundary (50% of the PPS modes and 40% of the PSS

modes) and within sediment boundaries (40% of the

PPS modes and 60% of the PSS modes).

The COT is found within a narrow 20-km zone,

which coincides well with previous studies in the area.

Low ray-coverage affects the mapping of the crystal-

line continental crust and the results presented here in

terms of composition are inconclusive. However, our

results suggest fractured metasediment in the upper

crust and an average felsic composition of perhaps

felsic amphibolite facies gneiss in the middle conti-

nental crust. The lower continental crust is only

mapped in the westernmost part near the COT leaving

Fig. 12. Geological model of the profile. The crystalline oceanic crust is believed to consist of alternating segments of normal oceanic crust and

mantle-derived crust (ultramafic rocks). The presence of ultramafic rocks is based on the averaged distribution of the magma-starved segments

generated at the ridge axis assuming a constant spreading direction.


lower crustal composition and total crustal thickness

unresolved in this area. The gravity data does however

indicate a f 24-km-thick continental crust.

The anomalous thick oceanic crust of 6.7 km could

be an effect of large magma production in the axial

highs along the ridge axis. Our results show that it is

not straightforward to determine lithology and phys-

ical properties of the crystalline oceanic crust on the

basis of seismic properties. A model with alternate

segments of normal oceanic crust and mantle-derived

crust bound by the distribution of the magmatically

active zones (axial and off-axial highs) and magma-

starved areas in between is considered as the best

alternative, although this model is unable to account

for large local variation in crustal thickness expected

under these conditions. The observations presented

here do reflect the complexity and opaqueness of ultra

slow spreading ridges and clearly show why these

end-members of mid-ocean ridges systems need to be

further investigated.

Acknowledgements

We thank the crew of R/V Hakon Mosby and

the technical staff from the Institute of Solid Earth

Physics (now the Department of Earth Science),

University of Bergen and the Laboratory for Ocean

Bottom Seismology, Hokkaido University, for their

help and support during the acquisition of the data.

Funding by the Norwegian Petroleum Directorate

(OD) is gratefully acknowledged. We finally thank

the reviewers for their many good and helpful

suggestions.


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Documents

Crustal transect from the North Atlantic Knipovich Ridge to the Svalbard Margin west of Hornsund