JOURNAL
The
OF GEOPHYSICAL
continent-ocean
crustal
RESEARCH,
VOL. 99, NO. B5, PAGES 9117-9135, MAY
transition
across
the
10, 1994
southwest
Greenland margin
Deping Chian and Keith E. Louden
Depanmentof Oceanography,
DalhousieUniversity,Halflax,NovaScotia
Abstract. This paperexaminesthe completecrustaltransitionacrossthe nonvolcanic,southwest
Greenland
continental
marginof theLabradorSeausingwide-angleandcoincident
verticalincidence
seismicprofiles.Six oceanbottomseismometers
anda sonobuoy
recordP andS wave
firstandmultiplearrivalsfromthecrustanduppermantle,whichareanalyzedby two-dimensional
dynamicray tracingandone-dimensional
reflectivitymodeling.Theresultingseismicvelocity
modelrequiresthatthepreexisting
30-kinthickcontinental
crustis thinnedabruptlyto ~3 km
acrossthecontinentalslope,primarilyby removalof thelowercrust.Fartherseaward,thecrust
thickensto -6 km primarilythroughtheadditionof a high-velocity
(7.0-7.6km/s)layerin the
lowercrust. This lowercrustallayeris 4-5 km thick,extendsfor a horizontaldistanceof-80 km,
andis interpreted
aspartiallyserpentinized
uppermantle.It is overlainby a low-velocity(4.0-5.0
km/s),upperlayerwhichis interpreted
ashighlyfractureduppercontinental
crust.Ourmodel
suggests
fiat seafloorspreading
did notstartuntilchrons27-28, 13 Ma youngerthanpreviously
suggested.
Thisinterpretation
is supported
by two-dimensional
modelingof gravityandmagnetic
dataalongtherefractionline. Ourresultsareconsistent
with a simpleshearmechanism
for the
initialrifting,withtheSW Greenland
marginastheupperplate.However,a full characterization
of theriftingmechanism
mustawaitcomparison
with a seismicmodelfor theconjugatemargin,
east of Labrador.
Introduction
A major goal of studyingthe structureof Ariantic-type,
passivecontinentalmarginshas been to discriminatebetween
different rifting models,especiallybetweenpure [McKenzie,
1978] and simple [Wernicke, 1985;Lister et al. 1986] shear
mechanisms. A simple shear model invokes a low-angle
detachmentfault that cuts throughthe continentalcrust and
separatesthe resulting margins into asymmetric,upper and
lower plate conjugates.The upperplate marginrecordsmore
thinning in the lower crust and less syn-rift subsidencethan
does its conjugate,lower plate margin. This model has been
supportedby the existenceof an S-reflector,interpretedas the
detachmentfault, and a smaller amount of upper crustal
stretchingas comparedto total crustalthinningin the Bay of
Biscay [Le Pichon and Barbier, 1987] and Galicia Bank
[Boillot et al., 1989b, 1992]. On the other hand, the observed
overall symmetry of crustal thinning acrossthe restoredrift
zone for the conjugatemarginsof FlemishCap [Keen and de
Voogd,1988] andGobanSpur[Peddyet al., 1989] suggests
a
pure shear,symmetricextensionwith final breakupcloserto
one side of the rift zone [Keen et al., 1989]. Finally, it is
possible to interpret the rifting of some margins using a
mixtureof pure shearfor the lower crustand lithosphereand
simple shearfor the upper crust (e.g., Galicia Bank; Sibuet
[1992]).
Seismic refraction methodshave been used in defining
crustalaffinitieson rifted continentalmarginsbut mainly for
volcanic passive margins where the existenceof a high-
Copyright1994by the AmericanGeophysical
Union.
Papernumber93JB03404.
0148-0227/94/93JB-03404505.00
9117
velocity (7.2-7.5 km/s), lower crust likely represents
magmaticunderplatingformedduringthe initial rifting [e.g.,
Mutter et al., 1984;Holbrook et al., 1994]. The initial rifting
mechanism, however, is probably best preserved on
nonvolcanic margins, where the preexisting velocity
signature of the continental crust has not been altered by
subsequentvolcanic activity. For investigatingthe nature of
rifting mechanisms,refraction methods must delineate the
detailed velocity structure from unstretched continent to
normal oceaniccrust. An example is the Goban Spur margin
where recent refraction results suggesta uniform, gradual
thinning of both the upper and lower crust (i.e., pure shear)
across the margin [Horsefield et al., 1994], which is in
agreementwith reflectiondata for the conjugatemargins[Keen
et al., 1989; Peddy et al., 1989]. On the other hand, the wide
zone of thin crust off the Iberia margin bears a velocity
structure
which
is different
from
normal
oceanic
crust and
therefore is assumed to represent thinned continental crust
[Ginzburg et al., 1985; Whirmarshet al., 1986, 1990]. The
existenceof a thin lower crustallayer bearinga high velocity
of 7.2-7.6 km/s in nonvolcanicmarginsis interpretedto be
either underplatedmelt [Reid and Keen, 1990; Whitmarsh et
al., 1990] or serpentinizedmantle [Pinheiro et al., 1992;
Reid, 1992].
The conjugatemarginsof the LabradorSea are ideal for the
investigation of passive margin formation. Reconstructed
positionsof the margins can be well constrainedfrom the
seafloor spreading history of the basin, given its small
overall width [Srivastava, 1978; Roest and Srivastava, 1989].
The margins are nonvolcanic and exhibit an apparent
asymmetry in the shelf/slope morphology, sediment
thickness [Balkwill, 1987; Rolle, 1985], and gravity
anomalies[Woodside,1989]. In thispaperwe presentdetailed
seismic velocity variations and their lithologic and tectonic
interpretationfor the complete transition from continent to
9118
CH[AN AND LOUDEN: CONTINENT-OCEAN CRUSTAL TRANSITION
oceancrustacrossthe SW Greenlandmarginof the Labrador extension
of the Canada-Greenland
continentpossiblystarted
Sea (Figure1). P andS wavesfrom denseairgunshots at ~160Ma, asevidenced
fromthedatingof coastparalleldike
penetratingcrust and upper mantle are observedon ocean swarms
in SW Greenland
[Watt,1969]. Finalriftingoccurred
bottomseismometers
alonga 230-kmprofile.Thesedataare with minimal volcanicactivity,leavingless than 800 m of
usedto constructa two-dimensional
(2-D) seismicmodel. This syn-riftextrusivebasaltson the Labradormargin[Balkwill,
model contains a 50-80-km wide transitional zone with a low-
1987]. Features
typicalfor Atlantic-type
passive
marginsare
velocityupperlayer interpreted
as thinneduppercontinental present,includingasymmetricshelf/slopemorphologyand
crust and a high-velocity lower layer interpreted as sediment distribution. Less than 4 km of sediments are
serpentinized
mantle.Our datasuggest
thatchrons27-28 mark deposited
nearthe westGreenlandmarginwhile the Labrador
theinitiationof seafloorspreading
in theLabradorSeaandare marginhas accumulatedup to 11 km of sedimentsunderthe
compatible with a simple shear mechanismfor the initial
shelf [Tucholke, 1988].
rifting.More detailedconstraints
on the riftingmechanism, The seafloorspreadinghistory of the LabradorSea is
however,
mustawaita seismic
modelfortheconjugate
margin, documentedby magnetic anomaly lineations, with welleast of Labrador.
defined anomalies observed between chrons 21 and 27
RegionalGeologyand Tectonics
linearionsolder than citron 24-25 are brokenby several
[Srivastava,1978;Roestand Srivastava,1989]. Anomaly
The Labrador Sea forms a northwestwardextensionof the
North Atlantic Ocean,with the main basinabout900 km wide
fracture zones (Figure 1), which are associatedwith a NW
rotationof the spreadingaxis at anomaly24-25 whenEurasia
separated
from Greenland.The lessclearlydefinedandlower
bounded
in thewestbyLabrador,
in theeastby SWGreenland,amplitude anomaliesin areas older than chron 27 have been
and in the northby Davis Strait(Figure1). The rift-related interpretedby Roest and Srivastava[1989] as chrons31 and
65
6O
5O
4O
6O
55
31
27
26,
'
24
i
60
r .......................
r
'
55
21
50
Figure1. Location
mapof theLabrador
Sea,including
bathymetric
contours
plotted
every1,000m (dotted
lines),numbered
magnetic
anomalies
(thinlines),fracture
zones
(thickgreylines),andanextinct
rift axis
(thick
broken
lines).
Refraction
lines
(thick
lines)
include
R2 (this
paper),
R1 [Chian
andLouden,
1992],
and
twolinesalongandacross
theextinct
ridgeaxis[OslerandLouden,
1992;Osler,1993].Labeled
crosses
are
OBSinstrument
locations
A-G(Table
1);circle
issonobuoy
F [Stergiopolous,
1984].
R/VMINNA74gravity
andmagnetic
profile
isnearly
coincident
withlineR2.
CHIAN
AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL
TRANSITION
9119
33. This has been challengedrecently by Chalmers [1991]
who, on the basis of reprocessingof multichannelreflection
data BGR-21 (the FederalInstitute for Geosciencesand Natural
Resources),suggestedan initiation of seafloor spreadingat
anomaly27 (62 Ma). His interpretationplacesthe continentocean boundary (COB) -120 km farther seaward of the
previously suggestedposition.
Recent refraction studies along the shelf of the SW
[1983] (Figure 2). In order to convert two-way travel times
into depth, we measured the P wave velocities for each
sediment layer from sediment-corrected ocean bottom
seismometers(OBS) profiles, with reference to the onedimensional (l-D) modeling results from the sonobuoydata
aroundODP Site 646 [Srivastavaet al., 1989]. A It value of
2.2-3.0 is obtained in transition zone II which is slighfiy
higher than the It value of 2.0 in zone I. The amountof
Greenlandmargin [Chian and Louden,1992] reveal a typical stretchingdeterminedin this manneris typically less than the
30-kin thick, preexisting continental crust without any stretchingof the crustas a whole. It is likely to representonly
magmaticmodification. In contrast,500 km north of this area the extensionin the most brittle layers of the upper crust [Le
on the Greenland margin of the northern Labrador Sea, Pichon and Barbier, 1987]. The remainingextensionmay be
refractiondata show the existenceof a 5-6-km thick, high- accommodated
by creepin the lower crustfor depth-dependent
velocity (-7.5 km/s) lower crust, interpreted as magmatic pure shear models or low-angle faulting for simple shear
underplatinginto the lower continentalcrustduring the initial models.
rifting [Gohl and Smithson, 1993]. It is likely that this
underplatingis related to hotspotactivity in Davis Strait at
Wide-Angle SeismicData and Modeling
chron25, by which time the centraland southernLabradorSea
was alreadyopened. There are scattered,old refractionlineson Data
the northernLabradormargin [Van der Linden, 1975] and on
SevenOBSweredeployed
alonglineR2, sixof them(OBS
the SW Greenlandmargin[Stergiopolous,1984] to controlthe B, C, D, F, and G) giving useful information. The OBS used
deep crustal structure for the Labrador basin. These data,
alongline R2 recordseismicsignalsfrom hydrophone
and
however, do not yield sufficient resolution, and additional vertical and horizontal4.5 Hz geophoneson cassettetapes.
deepseismicdata are requiredto furtherassessthe mechanism These analog data were digitized and interpolatedto a fixed
for the initial rifting.
sampling rate of 80 samples per second. Subsequent
Reflection
Profile
The seismicdata we presentincludewide-anglerefraction
and coincidentsingle-channelreflectionprofiles along a 230-
km line (R2) that covers the completetransitionfrom
preexisting continental to oceanic crust across the SW
Greenlandmargin (Figure 1). The seismicsourcefor bothdata
setsisanarrayof six,1000in3(16.41)airguns
firingevery
processingincludesband-passfiltering (4-9 Hz for P waves
and2-8 Hz for S waves),clippingof high amplitudes,andgain
that increaseslinearly with horizontal distance.
Since different recording channels sometimes respond
differently to various seismicphases,we find that a sum of
several channelsis a useful way to presentseismogramsthat
demonstrateall prominentphaseswith just one figure. Before
the summation, however, careful correction for the relative
time differencebetweenchannelsis necessaryto maintain or
enhancethe coherencyof major seismicphases.In addition, a
coherency filter is applied to some of the profiles. This
algorithm [Chian and Louden, 1992] is definedas a stackof a
number of traces along an optimal phase velocity, which
produces the maximum semblance within a moving time
window of 0.15 s, followed by weightingwith the normalized
semblance(0-1). The coherencyFilteringthus deemedcan be
usedto enhancecertainphases,particularlyS waves,which are
frequentlyentangledwith P wave reverberations
and multiples
(Figure 3).
Major constraints on the crustal velocity structure come
to boundary
R3/4at Ocean
DrillingProgram
(ODP)Site646 from OBS B, C, and D, which providethe best data quality.
[Srivastava et al., 1989]. The layer beneaththis boundary
The data from OBS E are of poorer quality, possiblydue to
containsthe syn-rift sediments,which are adjacentto the shelf
increasednoisefrom deepwatercurrents.OBS F andG, situated
and characterized as strongly folded and faulted seismic
undershallowwater (-200 m; Table 1), alsogive good seismic
sequences
(Figure 2).
The most significant feature of the profile is the signals, but the seismic rays are severely distorted by the
topographyof the basement. In general, the profile shows sharpshelf break and/or the Moho relief, making it difficult to
extract crustal information. Because of the failure of OBS A,
threezonesof distinctbasementtopography,separatedby two
we used data from an older Sohobuoyprofile [Stergiopolous,
basementhighs. From the shelf seaward,zone I with large
1984] to help constrainthe oceaniccrustalstructure.
tilted basement blocks probably represents stretched
continentalcrust. At the seawardend of the line (zone 111),
Reflectivity
Modeling of
where magnetic anomalies 26 and 27 are located, reflection One-Dimensional
hyperbola from the basement are typical of oceanic crust. Sohobuoy F
Zone II in betweenis characterizedby a numberof smaller
The wide-angle profile from sonobuoyF (Figure 4), shot
tilted blocks and is interpretedas transitionalcrust.
along strike during Hudson 79-013 (Figure 1), is used to
As a First-orderapproximation,the stretchingfactor for the constrainthe crustal velocity structureat the western end of
90 s with an approximatespacingof 200 m. The reflection
data were recorded primarily for constrainingthe sediment
thicknessand basementmorphology.
The pattern of reflectivity on the single-channelprofile
(Figure 2) enablesa grossdefinitionof threeprimary seismic
sequenceswithin the sedimentarysection. The first and the
secondlayersare characterized
by high-amplitudereflectorson
top and low-amplitudereflectorsbelow.A groupof continuous
seismicphasesforms the boundarybetweenthe secondand
third (bottom) layers. This boundary can be traced
continuouslyfrom the margin to the eastand may correspond
uppercrust• is calculatedbasedon the geometryof the tilted
line R2. An additionalsonobuoy,
but with poorerdataquality
blocksin zonesI and II following the methodof Chenetet al.
was also deployed at the opposite end of the same profile.
9120
CHIAN AND LOUDEN:
,
- .......
..
CONTINENT-OCEAN
.-,'-,:-;:-;
CRUSTAL TRANSITION
.....
, ....................
,
,
4
..........
. ................
i:........
•.:'.•:'.':'...,.
....''."";.:::: •.':..:'<........
:."."•'
•d•........
•,•.......
,•
•.•;•f•.,•
;:,•';:;;:::'-.,:::
.............
•.•8
:::::::::::::::::::::::::::::
,•.,..•:
.....:.•:.-,• ........: ,:,,...:-•
............
•............
:::":::::
• •.•.
::.:":::: ....
, .•:• .,. ••
.....
.• .•.. ,. .... •. .
200
•es foHow•g Roest a•
Sriv•t•a
[1989].
Previous slope-interceptanalysisof the reversedtravel times
[Stergiopolous, 1984] yielded similar velocity models,
indicating a homogeneousstructure along the line. In our
reanalysis of SohobuoyF, a full waveform record section
(Figure 4a) was producedby digitizingthe original shipboard
FM tape recording. Horizontal ranges were calculatedfrom
director reflectedwater wavesusingan averagewatervelocity
of 1.486 km/s. Determinationof the P and S wave vdocitydepth struc•e is basedon forward modelingof the observed
record section using 1-D ray tracing and full reflectivity
syntheticseismograms
(Figure4b). Completevelocity-depth
values
for this modelare givenin Table 2.
Crustal P phasescontain two separateamplitudemaxima
centered at 13-16 and 17-23 km, with relatively abrupt
changesin phasevelocity. To fit this patternrequiresa blocky
structure for the crust, with thicker units of low-velocity
S waves show a similar although less well-developed
patternof amplitudesand phasevelocities.The one exception
is thattheuppermost
crustal
layer(withVp=5.5km/s)must
represent
a gradient
zone(withVs=3.0to 3.'2kin/s),primarily
in orderto fit the S wave travel times.The two deeperconstant
velocitylayersare modeledwith Vs of 3.6 and4.1 km/s,
giving Poisson's
ratiosof 0.28 and 0.26 for the mid and lower
crust, respectively.
Figure 4b comparesvelocity models of SohobuoyF with
resultsof 2-D syntheticmodelingfor a separateprofile [Osier
and Louden, 1992; Osier 1993] which terminates70 km to the
southwest,as located in Figure 1. This comparisonindicates
that at sonobuoyF: (1) The uppercrustis thickerandcontains
a morecomplexstructure,and(2) The lowercrustis of similar
thicknessbut containsa smaller velocity gradient.
gradientat velocitiesof 5.5, 6.5, and 7.2 km/s, separatedby
higher gradient zones with variable thicknesses.The initial Two-Dimensional Modeling of OBS Data
velocitygradientfrom 4.8 to 5.5 km/s is not well constrained,
Since significantlateral variationsin seismicvelocity are
apartfrom the total travel time, on accountof the interference expectedacrossthe continent-oceantransition,we use 2-D
of the water reflection with the ground wave. The linear dynamicray tracing [Cervenyet al., 1977] for modelingthe
gradient
of 1.3s-1 from5.5 to 6.5km/screates
theinitial observedOBS data. The sedimentarystruc•e and basement
amplitudebuildupin the refractedwave (P2), althoughat relief used in the modeling are derived from the coincident
ranges slightly less than observed.The sharp discontinuity reflection data (see previoussection).We tried two methods
from 6.5 to 7.2 km/s createsthe second-amplitude
maximum for describingthe crustalstructure.The first one divided the
from the wide-anglereflection(P3P), although
it is not as crust into several homogeneous layers using velocities
sections.The depth of each
abrupt or as strong as observed. The depth and velocity obtainedfrom sediment-corrected
boundarywas varied along the profile to match the observed
transition at the base of the crust is poorly constrainedby
modelinga third set of phasesas PmP with increasing
travel
times
of crustal
and mantle
arrivals.
This
method
amplitudesat ranges27-30 kin.
generatedstrongwide-anglereflectedenergywhich doesnot
CHIAN
11 -
AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL
TRANSITION
OBSC
•
9121
Raw
9
11
Coheren½
Filtered
3
1
o
4
12
6
-2o
24
;
-32
i
40
44
Distance(km)
Figure 3. (top) Sediment-corrected
vertical geophonechannelfrom OBS C, band-passf'fitered(2-8 Hz) with
gain increasinglinearly with distance. (bottom) Samedata with coherencymixing of five adjacenttraces.The
mixing velocity is scannedbetween2.5 and 4.5 km/s at a stepof 0.5 lcm/s.
producean acceptablefit to the observedamplitudes.Small
misfits in travel times also existed. A second approach
thereforerepresentedmost structuralvariationsby lateral and
horizontalvelocity gradients.Velocitieson a regulargrid are
obtained by bilinear interpolation of values at arbitrary
points.
The seismogramsgeneratedfrom our final, 2-D velocity
model with this secondmethod are comparedto individual
recordedsectionsfor OBS B, C, D, E, andG in Figures4-8. The
strong seismicphasesthat are crucial in determiningthe P
velocity model include refractionsfrom the sedimentlayers
(Psed),refractions
fromtheupperandthelowercrust(P2,P3),
andreflections
andrefractions
fromtheuppermantle(ProP,
Table 1. OBS Positions and Water Depths Along
Refraction
LineR2
OBS
Latitude,øN
59.8408
60.0517
60.2767
60.4442
60.5516
60.6633
60.7400
Its curvaturecan constrainthe velocity of the lower crust and
its travel time the depth of the Moho when the first arrival
cannot be recognized (e.g., OBS B; Figure 5). Another
prominentmultiple(PmP') is produced
betweenthe seafloor
and the sediment/basement interface. This multiple is
strongestin the hydrophonechannelof both OBS B (Figure 5)
and OBS C but is minimal, for example, in the horizontal
geophoneof OBS C (Figure6).
OBS B. The velocity-depthstructureof crust from the model
of sonobuoyF (Figure 4) is used as a control for the western
end of the 2-D model. The upper crustalstructureis smoothed
laterally to provide an overall fit to the westernprofile of OBS
B (Figure 5). The velocity (7.15-7.25 km/s) of the lower crust
and the depth (12 km) of the Moho determinedfrom modeling
the curvature
of ProP2to thewestof OBSB aresimilarto the
Longitude,
øW
Water
Depth,m
A
B
C
D
E
F
13
Pn)' In addition,two setsof strongmultiplesare sometimes
useful.Thewatercolumn
multiple(PmP2)is frequently
strong.
52.1033
51.2900
50.4233
49.7767
49.3514
48.9100
48.6133
3368
3303
307 6
2916
2823
136
175
model of sonobuoyF. The lower crustalvelocity eastof OBS
B is less well constrainedbecauseof distortioncausedby the
suddenbasementjump.
OBS C. The seismograms of OBS C (Figure 6) are
characterizedby significantdifferenceson both sides of the
OBS. On the westernside, the velocity of the upper crust is
about4.0-5.0 km/s. This low vdocity forms a strongvelocity
contrastwith the layersbelow, producinga distinctreflection
phase(P2P). The lowercrustis represented
by a clearphase
P3, whichis refractedby a gradientlayer of high velocity
(7.0-7.6 km/s), with the bulk of the lower crust at 7.4-7.6
9122
CHIANAND LOUDF_•:CONTINENT-OCEAN
CRUSTALTRANSITION
S/B F
Filtered
10
4
•
0
2
4
6
10
20
30
10
20
30
8
•,.,
ß
.
!
4
0
Distance(km)
Figure4. (top)Sonobuoy
F hydrophone
filtered
at4-10Hz.Overlain
linesaretravel
timecurves
computed
fromoneLdimensional
reflectivity
modeling.
Ranges
arecalculated
fromdirect
water
wave
(notshown)
at0-10
kmandfromwater
bottom
reflection
at10-30
km,assuming
anaverage
water
velocity
of1.486
km/s.
(bottom)
One-dimensional
reflectivity
synthetic
seismograms
generated
fromtheP andSvelocity
model
(solid
line)in
theupperleftcorner.
Complete
velocity-depth
values
forthismodelaregivenin Table2. Dotted
linein the
inset
shows
theP velocity
structure
70kmsouthwest
ofSohobuoy
F [Osler,
1993].
P2and
S2 arerefractions
from
upper
crest;
P3P and
S3Sarereflections
from
lower
crust;
Propand
SInS
arereflections
from
upper
mantle.
S wavesaredoublyconverted
fromandtoP wavesat thesediment-basement
interface.
km/s.On the eastern
sideof OBSC, thereis a significantof OBS D (Figure7) at westernrangesof 20-28 km. The
traveltimepull-down(by 1 s) of uppercrustalarrivalsbetween modeling
of Pn for thewestern
ranges
of theOBSsupports
the
ranges
of 12 and20 km whichcannot
befit unless
a velocity velocitystructures
underthebasement
high as obtainedfrom
of -6.5 km/s is introducedbeneaththe elevatedbasement.The OBSC. Ontheeastern
sideof OBSD, theuppercrustal
arrival
highamplitudes
of thelowercrustal
arrivalP3 (at easternis well modeledandcorresponds
to an easterlyincreasing
ranges
of-30 km) areproduced
by thelargevelocitygradient velocityof 5.3-5.6km/s. Firstarrivalsat ranges>28 km have
under this elevated basement. The lower crust under this traveltimestooearlyto beproduced
solelyby shallowing
the
basement
highhasa velocity
of 6.8-7.4km/sconstrained
by basement toward the continent. A reasonable fit can be
P3 witha phasevelocityof-7.0 km/s.A highervelocity produced
by settingtheMohoat a veryshallowdepth(-8.2
would
incorrectly
reduce
thetravel
timeofP3-TheMohonear km) andthemantleunderneath
with a highvelocitygradient
thebasement
highis placedat a depthof 11.8km whichis the (7.7-8.0 km/s). It should be noted that the calculatedtravel
sameas that westward.This providesa reasonablefit to the
times are slightly earlier than observed.While a smaller
traveltimeof PrnP whoseslopeis distorted
by thefaulted velocity of 7.5-8.0 km/s producesa better fit, this would
basement
blockeastof thebasement
high.A strong
PmP
degrade both the calculated travel times for OBS E and the
phaseis notobserved
westof OBSC butis replaced
by the computedgravityanomaly.Thereforecomplications
existfor
clearappearance
of Pn whichcanbe tracedto a horizontal this part of the seismicmodelwhichour dataare not sufficient
rangeof 100 km. The steeptiltingof themantlearrivalat a
western
rangeof 40 km is caused
by thebasement
jump(>1
km) adjacentto OBS B.
to resolve.
OBS E. Shotsto the westof OBS E (Figure8) give
importantinformation
on the depthof the Moho.Between
OBS D. The highvelocityof the elevatedbasement
is also rangesof 22-30km, a strongphasecanonlybe modeled
asa
expressed
in the traveltimepull-downof lowercrustalarrivals refraction
froma layerwitha velocityof >7.7km/sat 8-9 km
CHIAN AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL TRANSITION
OBS B
Hydrophone
10 ....,,.•.,;,. .......'............
,4",'""'.
......,,,.•.........
,..,'.....
;-,•;._'-.•;.,,..;;"•;•,;
...........,,
.........
;;;,'.,.4.'..,,.;,,v,,.,,.
........
,.• ....;;';,'•
,• ,0"•'
',4;"'.•'.
,'"•;';."".",,,
"•......
"•'•"•"'
• ...... ,........
,.......
•. •,.::,,=
.........
•
.....
:..........
,.•,,..,,,,,
........
.• .....
•,• "' '"",,,:,:.•,,,-•
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,,,,,,,,,
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..............• .....,',"•,:
.........
2 I'•:,:?•
:,•;::•,.:,,:::,::,',;::.,•:::'::'•['
,.•,:•'•,:?,:,,,:•,:,'::•::',,:'.'d,.,:•
50
40
30
20
10
0
10
20
30
40
50
2D Synthetic
10
•8
Distance(Km)
Figure 5. (top) OBS B hydrophone
coherency
filteredusingfive traces.Calculatedtraveltime curvesfrom
two-dimensional
dynamic
raytracingaresuperimposed.
PmPis thewide-angle
reflection
fromtheMoho.Prop2
andPmP' are water and sedimentmultiplesof PmP. (bottom)Syntheticseismograms
and travel time curves
generated
by two-dimensional
dynamicray tracingfromthe seismicmodelof Figure10a.
depth, which is consistentwith the modelingof OBS D. A
slower velocity will causean unacceptabletravel time delay
and will shift the major amplitudewestward.To the east, the
seismogramsare noisy and mostly affectedby basementrelief
acrossthe continentalslope.
OBS G. OBS G (Figure 9) is situatedat the easternend of
refraction
lineR2. P andS wavesfromtheuppercontinental
crust are observedwhich give the samevelocitiesas those
from anotherrefractionline R 1 alongstrike[Chian and
Louclen,1992]. Deeperarrivalsare distortedby the sharpshelf
break andhencegive little informationaboutcrustalstructure.
The f'mal velocity model that producesall the synthetic
seismogramsand travel times in Figures 4-9 is shown in
Figure10a. To showthe ray coveragefor differentregionsof
the model,we plot in Figure 10b the boundaries
of the model
overlainwith ray pathsfromOBS B, C, D, andE.
Determining
$ Waves
We presentthe f'trstshearwave data for transitionalcrust
anduppermantleacrossa nonvolcanic
riftedmargin.Thereare
primarily two typesof crustalS wavesobservedin our OBS
data: PSP and PSS (Figure 11). PSP phasesare formed by
double P-S conversion at the sedimenffbasementboundary.
Table
2.
One-Dimensional Velocity Model for
SonobuoyF
Thickness,
km
rs,,krn/s
3335
1.485
0.10
0.58
1.21
0.10
0.20
0.40
0.50
0.15
0.15
0.15
0.15
0.15
1.20
0.10
1.80
1.80
1.90
2.10
2.50
4.80
5.10
5.50
5.50
5.60
5.80
5.95
6.10
6.30
6.45
6.80
7.15
7.25
8.00
vs
•?s
0.58
0.64
0.85
2.70
2.90
3.00
3.15
3.20
3.30
3.45
3.50
3.5 5
3.60
3.85
4.10
4.15
4.70
9124
CHIAN
AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL
TRANSITION
OBSC
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-90
-80
-70
-60
-50
-40
-30
-20
-10
0
10
20
30
40
50
30
40
50
2D Synthetic
2
-100
-90
-80
-70
-60
-50
-40
-30
-20
-10
0
10
20
Distance(km)
Figure 6. (top) OBS C horizontalgeophonewith tracesbetweenrangesof-100 to 40 km coherencyfiltered
usingfive traces.Calculatedtravel time curvesfrom two-dimensional
dynamicray tracingare superimposed.
(bottom) Syntheticseismograms
and travel time curvesgeneratedby two-dimensional
dynamicray tracing
from the seismicmodel of Figure 10a.
The other prominentphase is PSS, which is only converted upper mantle, they are 1.83 and 1.90, respectively. This
once. The energy of such convertedS waves is especially methoddoesnot result in a totally satisfactoryfit to the data,
strongwhentheP velocityin thebottomof sediments
(Vsed) but it does provide a simple and reasonable estimate of
is closeto theS velocity(Vs)in thecrust[Spudich
andOrcutt, Poisson's ratio for each layer with error estimates (see
1980b].
Appendix). One advantage of this method is that all the
S waves can appearon all componentsensorsof the OBS. structural variations from the better observed P waves are
For example, the PSP phase at OBS B only appearson the retained for the modeling of S phases,which in generalare
horizontal geophone(Figure 11a), whereasthe PSS phasein less completeand def'mifive.
OBS C appearson both hydrophoneand vertical geophone
channels(Figure lib). Reverberations
of P phasesfrequently Error Analysis
interfere with S waves, requiring coherency mixing to
Since the final velocity model includesstronglateral and
suppressphaseswith higher apparentvelocities.In addition, vertical variationswith variable seismicray coverage(Figure
when sedimentphasesare strong,they interfere with upper 10b), analysis of the resolution of the model can only be
crustalS wavesbecauseof their similar velocities,making it investigatedfor one crustalunit at a time. We perform error
analysis for the lower crust but not for the upper crust
more difficult to pick the correctS phases.
Once S phasesare identified,their travel timesare modeled primarilybecausethe uppercrustdoesnot have sufficientray
by adjustingthe V_/V. ratio for all layersof the P velocity coverage.Reflectionsand refractionsare treated separately.
model.For OBS B, a V.JV. ratio of 1.8 for boththe upperand For a refractedphase,the velocity and top interfacedepthof a
lower crustenablesa fit of travel timesfor PSPandPSS (Figure layer are variedto determinetheir affecton the averagetravel
11a). The modelingof the PSS phaseat OBS C requiresan time misfit. For a wide-angle reflected phase, the velocity
additional
unknown:
thevalueof Vp/Vs ratioforthetop gradientandthe depthof the bottominterfaceof the layer are
sedimentlayer. This is determinedby maximizing the overall adjusted.When normalizedto the misfit endurance(maximum
these
become
Z2 plotswhere
traveltimefit. TheV_/Vs ratiosfor thethreesediment
layers traveltimemisfitacceptable),
are4.5,2.7and1.7.t'For
thehigh-velocity
lower
crust
and the bounds of
acceptableparmeters are delimited by the
CHIAN
AND LOUDEN:
CONTINENT-OCEAN
TRANSITION
.......,•,.,,,,
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9125
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;'•I,,,; .,' ';.,,::,',:,',::?
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2 [,•/•,;:;:,½,','
;.:,:,:'
,,.,,::::,',:,:,1,::.;:,
,.,',•:,
50
40
0
10
20
30
40
50
40
50
2D Synthetic
50
40
30
20
10
0
10
20
30
Distance(kin)
Figure7. (top)OBSD horizontal
geophone.
Calculated
traveltimecurves
fromtwo-dimensional
dynamic
ray
tracingare superimposed.
(bottom) Synthetic
seismograms
andtraveltime curvesgenerated
by twodimensional
dynamic
raytracingfromtheseismic
modelof Figure10a.
contour
of•2=1.Thedetails
of thismethod
arediscussed
in the spreadingrates of Roest and
Srivastava [1989] and the
time scale of Kent and Gradstein [1986]. The observed
FigureA1 givesresultsof an erroranalysisfor V_ of the magneticdata (Figure 12) show a clear changein patternat
lower crust west of OBS C. The contour levels are normalized
anomaly28. While chrons25 and 26, and possibly27 and
to 40 ms, which includes the errors in first-arrival travel 28, correlate well with observed anomalies, chrons 31-33 do
times.
Errors
are+0.10km/sforVpandi-0.3kmforthetop not. Thereforeit is possiblethat seafloorspreadingwith the
boundary
depth.Thesameanalysis
is performed
for thePn generationof true oceaniccrust did not start until chron 28.
the appendix.
phase,whichindicatesan errorof +0.05 km/s for the upper This is compatible with the interpretationof Chalmers
mantleP velocityand+1 km for the Moho depth.For a PSS [1991].
phaseat OBS C, erroranalysisfollowsa similarapproach
(see
The gravity map of the LabradorSea [Woodside, 1989]
Appendix) which indicates an error of :[-0.20 km/s for the
lower crustalS velocity(FigureA2).
showssmall along-strike
variationsadjacentto line R2,
justifying two-dimensional gravity modeling. The
The errors of P and S velocities for the lower crust west of
contributionof the sedimentsis calculatedusingthe drilling
OBS B areestimated
by iteratingthe2-D modelfor OBS B and resultsat ODP Site646 [Jarrardet al., 1989](seeFigure1 for
1-D reflectivitymodel for Sohobuoy
F. By this qualitative location) to determinethe velocity-densityrelationshipfor
approach,we estimatethatthe errorsare approximately
+0.15 the top 1 km of sediments.These data are extrapolated
km/s
forVpand:[-0.20
km/s
forVs.These
values
aresimilar
to
the resultsof the previous,morequantitative
approach.
Gravity and MagneticModeling
The gravityand magneticdatausedfor modelingwere
extractedfrom a R/V MINNA 74 profile[Voppelet al., 1988]
(S. P. Srivastava,personalcommunication,
1993), which is
downward
to
cover
the
entire
sediment
column.
The
contribution
from the crustis calculatedfrom the velocitygrid
obtained from the 2-D ray tracing model (Figure 10a),
convertedinto densityusingthe relationship p = -.6997 +
2.2302V..
-.598Vp
2+ .07036V.
3- .0028311V.?
[Ludwig
et
al.,1971•.
•'crustal
•'as2.40g/cm
Wesettheminimum
density
3
andthemaximum
as3.30gJcm
3.Weassume
auniform
mantle
(3.30gJcm
3) andcalculate
gravity
fromblocks
above
almostexactlycoincident
withourrefraction
lineR2 (Figure density
1). Calculationof syntheticmagneticanomaliesare basedon
35 km depth.
9126
CHIAN AND LOUDEN: CONTINENT-OCEAN
CRUSTAL TRANSITION
OBS E
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Figure8. (top) OBSE verticalgeophone
coherency
filteredusingfivetraces.
Calculated
traveltimecurves
from two-dimensionaldynamicray tracingare superimposed.
(bottom)Syntheticseismograms
and travel time
curvesgeneratedby two-dimensional
dynamicray tracingfrom the seismicmodelof Figure 10a.
It can be seen from Figure 12 that the gravity contours
follow velocity contoursexceptfor the leftmost50 km of the
profile. This exceptionis introducedbecausestrict adherence
to the velocity-densityrelationshipused in the remainderof
crustal structure result from this model:
the model producesa calculatedgravity curvefor the leftmost
1. The observedfree air gravityanomalyalongline R2 50 km which is --20 mGal higher than observed. An
provides the only control over the shape of the deepening alternative for avoiding this misfit is to deepen the Moho
Moho eastof OBS E. It is characterized
by a positivepeak (34 from 12 to 13 km west of OBS B. While this changeof Moho
mGals) at the sheffbreak,followedat 10 km seawardby a large is within the error boundsof our data, it is not supportedby
negative anomaly (-63 mGals). This sudden change is any otherdeepseismicdata in the basin.
primarily due to the edge effect betweenthe abruptdeepening
Discussion
of the basement and the steep slope of the Moho. The
calculatedgravity low between180 and 200 km is particularly The Seismic Model
sensitiveto the Moho relief under the continentalslope.
The seismic model shown in Figure 10a representsthe
2. The model providesimportant controlsover the depth of
complete
crustal transition across the SW Greenland
the Moho at the baseof the continentalslope.The diagramin
the middle of Figure 12 shows two computedgravity curves. continentalmargin.A grossfeatureof the model is that it can
The f'malcomputedgravity curvefits the overallpatternsof
the observeddata with a maximummisfit of 5 mGals (Figure
12), which is not significantconsideringthe resolutionof the
observed gravity profile. Two important controls on the
The
dotted
curve results
from
a Moho
which
remains
flat
seawardof the continentalslope. A Moho which shallowsto
8.2 km at the bottom of the continentalslope significantly
improves the fit of the gravity (dashed line) and is more
consistentwith our 2-D seismic modeling, particularly for
OBS E (Figure8).
be divided
into
three distinct
zones, similar
to those
previouslyidentifiedfrom basementmorphology.In zone I,
the 30-kin preexistingcontinentalcrustis thinnedabruptlyto
a thicknessof only -3 km under OBS E. Furtherwest, this
layer terminatesat a basementhigh where it joins zone II,
which has a low-velocity upper crust, underlain by an
CHIAN AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL TRANSITION
OBSG
9127
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60
50
40
30
20
10
0
2D Synthetic
•
-'
2
•--'
60
50
40
30
20
10
0
Distance (km)
Figure 9. (top) OBS G verticalgeophone
coherency
filteredusingfive traces.Calculatedtraveltime curves
fromtwo-dimensional
dynamicray tracingaresuperin,posed.
Pg andSg (withassociated
phasevelocitiesin
km/s)areP andS refractions
fromuppercontinental
crust. (bottom)Syntheticseismograms
andtraveltime
curvesgenerated
by two-dimensional
dynamicray tracingfromtheseismicmodelof Figure10a.
unusually high-velocity (-7.5 kin/s), 4-5-km thick lower
the upper crest in the transition zone is a strong reflective
crust.This zonespansa lengthof--50 km, terminating
in the
boundary.This reflectoris similarto theS reflectorin theBay
west at anotherbasementhigh which separates
zoneH from a
7-km thickcrestbearingoceanic-type
velocities(zonem).
of Biscay [Le Pichon and Barbier, 1987], and it can be traced
in the depthsandvelocities
nearthe two basement
highsthat
separatethe three zones. Among these complexities,the
100 km westwardfrom OBS E . Thereforethe upper crust
throughoutzoneH possiblystemsfrom a commonorigin.
The mostsurprisingresultof this work is the existenceof a
high-velocity (7.2-7.6 km/s), lower crust which is 4-5 km
structuresaroundthe basementhigh betweenzonesI andH are
thick within a 50-70 km-wide zone around OBS C. The bulk
Remarkable vertical and lateral variations are observed both
crust
hasa Vpof7.4-7.6
km/swithanerror
themost
interesting
andimportant.
Totheeast,
theVpofthe partofthislower
uppercrestrangesfrom -5.5 to -6.2 km/s. The similarityof
this value with that of the adjacent,preexisting,upperand
middle continentalcrustunder the shelf [Chian and Louden,
1992, Figure 12] suggeststhat the crest east of OBS D is
composedmainly of stretchedmetamorphosed
upper and
middle continental crust.
To the westof thebasement
high,the uppercrestin zoneH
has critical importancein determiningthe position of the
continent-ocean
boundary(COB). An evident featureof this
upper crust is that the seismic velocity (4.0-5.0 km/s) is
unusually low compared to both the continental crest to its
east and oceaniccrust to its west. Comparisonwith a recent
multichannelreflection profile (Keen et al., 1994), which is
adjacent
andparallelto ourlineR2, indicates
thatthebaseof
boundof +0.1km/s(FigureA1). Thecorresponding
Vs is 4.0
+ 0.2 km/s as determined
from a PSSphase(Figure11), which
implies a Poison'sratio of 0.30 + 0.03. This high-velocity
layer is restrictedin its extentbetweenOBS B andD. Its high
velocity (7.4-7.6 km/s) cannot be found either in the
continentalcrest to the east [Chian and Louden, 1992] or in
the oceaniccrest to the west (sonobuoyF, Figure 4; OBS B,
Figure 5).
To investigatethe originsof the lower crestin zonesH and
IN,a.Vp- Vsdiagram
(Figure
13)isused
tocompare
reported
seismicdata for the crest and upper manQe[Christensen,
1989] with laboratory data for serpentin,,es and
pyroxene/olivine gabbros [Spudich and Orcutt, 1980a;
Christensen and Sinewing, 1981]. In addition, we include
9128
CH]AN AND LOUDEN:
CONTINENT-OCEAN
2D RAY TRACING
West
S/B F
OBS
CRUSTAL TRANSITION
MODEL ACROSS SW GREENLAND
C
B
MARGIN
D
F
East
G
E
.,•
2.3
2.6
I
I
I
26
27
28
25
VerticalExageration3:1
30
35
5'0
160
1•0
260
230
Distance (km)
10
150
230
Distance(km)
Figure10.(a)Two-dimensional
velocity
model
showing
contours
of Vp(inkm/s)across
theSWGreenland
margin.The contourintervalis 0.2 km/s.At somesignificant
velocityboundaries
(thicklines)contourlines
are terminatedfor clarity. Solidstarsindicatepositionsof OBS; oval delineates
the approximate
positionof
sonobuoy
F. Verticalexaggeration
is 3' 1. (b) Selected
ray pathsfor OBS B, C, D, andE overlainby boundaries
of the two-dimensional
model shown above.
resultsfor the lower continentalcrustalongthe SW Greenland mantle [Christensen,1966]. Therefore the velocitiesalone are
melt
shelf determinedby Chian and Louden [1992]. The observed not able to definitelydistinguishbetweenan underplated
Vpofthelower
crust
inzone
m • similar
tovalues
reported(i.e., high-Mg mafic gabbro)and serpentinizeduppermantle
by Osler [1993], who showsthat rather high velocities(<7.3 minerologies.
km/s) for layer threeare typical of the oceaniccrustalstructure
Undercrusted Serpentinite or Magmatic
within the Labrador Sea to the east of the extinct rift axis. The
Underplating?
even higher V.. for the lower crustof zoneH lies outsidethe
High-velocity lower crust has been found on other
region
ofnormal
gabbroic
crust
butisconsistent
withtheVp
observedfor the underplatedlower cruston volcanicmarginS. nonvolcanic,rifted continentalmargins,but its lithologyand
It also intersectsthe theoreticalcurve (PSP) for serpentinized mechanismof formationare similarlynot clearly determined.
CHIAN AND LOUDEN: COUNT-OCEAN
CRUSTAL TRANSITION
OBS B' S Waves
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••%,,,•
•.
,....
-.,". I':....•'•::'•.... ,. ,'" ' "•'•;.....
'•33 ' •.'":;'..•••"•"
•__..•-
,. ,•,,•.,.,,••••••,,,,,
,.,'
' ':"":
....,•--";:,'"•2•%•""
,,,,,....., ......
•
' ;."•....•..... •:•'.•"":',,
,"':;"
', •'" • •", ' "'......', •,,:'T'•,••"
4 I•;•;?;:
5O
40
30
20
,,,,•
',:•
•-•,•
•......,,(•%,:.•,•
•
',,
I,..,•",,, ,"'
'•
..... ,•.•',
:-•..,.•.,
w,"-- -•
--•r
--•
,• ,.,,
•
....,,•.....
•
..........
• ",''"",""', ,•
,........
,•.m,,.-m,•
,•.•, ,•;•
.•,
•.,,,"
.....•,... ,, •.',,•.•
•,,•."
_..... •.".•':'••
,••
.... ,,,,,•,•
,,, , ....•, ,. •,,,,, , ,,,,,,, ;, ,,,,,, ,';,,......,,•,,,,
50
•,';,,:,
• ....... •.,•,,
••••.,,
•,•,,,•,.,',,,•"•,•1
,,',,,•,•,'. ,• •
40
'-•.,,•,•."•
.,.. ,,,, ,.,.... • .....
•
a)
Hydrophone+ Vert Geophone
......
•L, •'•,• .......,•,•.,,•,
.,r•.
....,,,......•,,........, • .....
• ....,,, ...........
; .......
,..................
,;•,; ......
30
•.•"'•••-•),ip•
i•',,'..;•,'•,•;;,,.,, •,,,'. -.•-•,,
'........
'......
""" ' .........
' .......
",,,'•),),•,,•
.,•."'
....
."•,'"'
•' '", •,,"
,,,,,'.•
'••h
,,• h,).
I,,,.
• ,,,,•'•)'•
•," .".,,•
•,•
,
,.,,••.,,,••;,,,,, ,• --•
.....'•.......
•''' .... •'•,•
.,;...........,•''' ,,,""
....
,.............
•.•
.......
.' ,
.....,,•
•"..........
''';:L"',,,
'"::'•
",,.':; "'• :'•''".........
;:.........
;'"",,,....
", '"'",,'•
:'"q
.... ',........
;,.....
' ,'"' •;;
[,..,,',,,•:,
,..•,,.....>,,:,,.•?
.....
*',•
.....
,,,:
•,••-,,:,
,,•,.'i,•-.•.,,• :..,•=•,,.•;•
,%,,,, ..• .....
,.
. .,,,, ,,,,,,,,..
. ',•;.•,
...........
,....................
.,•,,
...., ......
,,
OBS C: S Waves
10
,,
C''"• '•.... :• "',.'""'"',".,.•"'".•
.........
•"',,,•,.'•'.•"•,•'• .... ß........ ' ";"
..........
' ' '
•
40
•"• .•., ' ,•
, ,,;•• .....,, , , .....•.....•,,.,, ,,,,......,......
,.....
,.....
. ,., ..........,,..........
;,,•,
,,,,,,,;;,,,,,;,,:,,
,,,,,,,,,;;
........
;;:
...........
,,,,,,,,.,:;;,,,,,,,;,;
..........
;;,",;;,,,:;;',,,,,•:
....
•,,, .,
2
5O
9129
.... •, ,•.,•..•'
,•
.•,;,.,••.,.,,•
•
,, ,,,,
•',/•,,•,,•
• ....... ;•.......... ,......
,,,..
.....
•:•:,•,,,,,
..........,•
. •..--.,,•,,
.....
•,,,,,;
-,
"--"'".-•,.,•'•'•,,.••,:'•,,,,•:•
•,, .... '• .........
•,,"•,:•
,',•',•,;,.,'"'•?•
..... •,I .... ,
,, •',--•r.
....
',',:
......
"'•'•"••'"
•'••,,
• •.'•'"'
,',,•.......
.• .......
"•,• '•','
•".•'" ""'"""•• ' '.'•I'•'
..... .,,,- •- ,,••••.,•••,,,,•.•,,.•,,•
,,•,,," •"Z•, ",•,'
•••""1
•..,,•,• _.•
,.• .......•, ,.•,, .....,, • ', ,,• ........
'.•, ".• •,,,' ',.•,, ',';
. ',"'r,,'•',,,'.... ,r',',,'
•,'•'•
• ....I,,;',;'•
•,'',•",,'•,,,"',',',',:
........
"', •',•' ', ','"•':,,•',"•,•:,•••
10
0
10
2O
3O
4O
50
Distance(km)
Figure 11. (a) CrustalS phasesfor the horizontalgeophone
channelof OBS B and (b) a combination
of
hydrophone
andverticalgeophone
channels
of OBSC. PSPis a crustalS refraction
generated
by doubleP-S
conversion
at the sediment/basement
interface.PSSis converted
onlyonceat thisinterface.
For example,on the westernmarginof Iberia the lower crust seaward-dippingreflectors), a feature absentfrom reflection
near the proposedCOB has anomalously
high velocity(7.3- profiles across the margins of the Labrador Sea [Balkwill,
7.6 km/s) and a small thickness(2-3 km). This has been 1987; Keen et al., 1994]. In addition, a melt model cannot
interpreted
eitheras underplated
melt [Whitmarshet al., 1990] explainwhy our observedhigh-velocitylayer is offsetseaward
or as serpentinite[Boillot et al., 1992]. Farther southin the by 80 km from the continentalslope,where underplatingis
TagusAbyssalPlain, even highervelocities(7.6-7.9 km/s) generally observed.
for thelowercrustareinterpreted
asserpentinites
[Pinheiroet
For the abovereasons,we suggestthat the observedhighal., 1992]. On the conjugatemargin off Newfoundland, velocity lower crust more likely representsa serpentinized
refractiondataon thesouthern
GrandBanksshowa thin,high- manfie. The observed V_b' - V.•' for the lower crust of zone 11
velocity(7.2-7.7km/s),lowercrustinterpreted
as serpentinite (Figure 13) is consistentwith the theoreticalcurve for-15%
[Reid, 1994]. A similarhigh-velocity,lower crustis observed serpentinization. Above the seismic Moho, the modified
on thesouthern
[ToddandReid, 1989]andnorthernmarginsof manfie has been "undercrusted"to form the high-velocity
FlemishCap [Reid and Keen, 1990],but it is interpretedas lower crust [Boillot et al., 1989a,b]. A small amount of
underplatedmelt.
serpentinization
alsopossiblyexistswithinthe uppermanfie
Amongthe two possibleexplanationsfor the origin of the layer, increasingits verticalvelocitygradientandproducinga
high-velocitylayer, the melt modelis favoredby the widely strongrefractedarrivalPn (Figure6). The basement
high,
diaper.It
acceptedexistenceof magmatic underplatingimmediately borderingzonesI and1I, may representa serpentinite
underneath the stretched lower continental crust of volcanic
lies on top of the most highly serpenfinizedlower crustal
margins [White and McKenzie, 1989]. However, for a
layer where the velocity is lower than on either side (Figure
crystallizedmelt to have a velocity as high as 7.5 km/s, a 10a). Compared with the serpentinizedridge off Galicia
high asthenospheric
temperatureis required, which for a margin [Boillot et al., 1992], however,greater along-strike
stretchingfactor of >5 would .haveproduceda minimm
variationsprobablyexist for the serpentinitediapir off SW
thicknessof 20 km. Sucha large thicknessis not observed.A
Greenland,as evidencedby its reducedelevationon a nearby
melt model also would require extrusivevolcanics(e.g., deepmultichannelreflectionprofile (Keen et al., 1994).
9130
CI-t[AN AND LOUDEN:
600f
CONTINENT-OCEAN
CRUSTAL TRANSITION
Computed
Magnetic
Data
215 •6....... Observed
Magnetic
Data
E 200
..........
ß
Eo
('9-200
-400
0
b
50
'
100
150
./
• ........
200
50
25
...............
-....
-..........
.....
-25
-- -- Observed
F.A.
Distance
inkm
Anomaly
'"-m
....................
Below
o•
'•
-;
200
.....
'""..'"'.• I
Computed
Gravity
Profii•
".\ ..'4,!
-50
Oceanic II
....
Transitional
: -•
"
:,...•,
.,...
Continental Depth(km)
,5
lO
vertical
Exageration
MANTLE
•,•-'"---2.83•--•
15
2O
25
30
Figure 12. The modelingof magneticand gravity data acrossthe SW Greenlandcontinentmargin. (a)
Observedmagneticdata (solidline) alongwith computedcurve(dottedline). The spreadingrate and direction
are from Roest and Srivastava[1989] usingthe time scaleof Kent and Gradstein[1986]. (b) Observedfree air
anomaly (solid line) along with computedgravity profile (dashedline) calculatedfrom the two-dimensional
densitymodelshownin (c). Dottedline is computedanomalyassuminga flat Moho at a depthof 11.8 km. (c)
Two-dimensionaldensitymodel. The density in the crust is calculatedfrom the two-dimensionalseismic
model
(Figure
10a)according
totherelationship
p =-.6997+ 2.2302V-.598V2
+ .07036V
3- .0028311V
4
[Ludwiget al., 1971]. Gravity contoursfollow velocitycontoursexceptfor the leftmost50 km.
Processes of
Crustal
Thinning
of a reflectoracrossthe SW Greenlandshelfin a nearbydeep
multichannelreflectionprofile (Keenet al., 1994).
If the above interpretations are correct, the lower
continentalcrusthasbeenremovedfrom this marginandmust
The major characteristics
of the uppercrustin zone II are its
presentlyexist under the conjugatemargin of Labrador.In the
low velocity (4.0-5.0 km/s), its small thickness(-2.5 km),
its blocky faulted pattern, and the irregular magnetic most basic form, this result suggestsa simpleshearmodel for
anomaliesit generates(Figure 12). It cannotbe oceaniccrust the initial rifting [Wernicke, 1985; Lister et al., 1986;
(i.e., extruded mantle melt) if our interpretation of the Wernickeand Tilke, 1989], in which the SW Greenlandmargin
acts as an upperplate and there is no associated
magmatism.
underlyinglayer as serpentinizedperidotite(as opposedto a
plutonic intrusion)is correct.We suggestthat the upper crust The very low angle normal fault that was responsiblefor the
in zoneII is composedof thinnedandfaulteduppercontinental final breakupmight have followedthe presentcurvatureof the
crust. An approximatecalculationof the stretchingfactor •
Moho under the continental slope, turning almost
based on the surface geometry of the tilted blocks suggests horizontally at the base of the thinned continental crust for
[•=2-3 for this part of uppercrust(Figure2). Fromthisvalue, about 100 km (Figure 12). We do not know if the original
the thicknessof the preexistingcontinentalcrustis estimated lower continentalcrustbeneaththis fault was affectedby synto be 5-8 krn. This is compatiblewith refractiondataalongthe rift ductile creep, in additionto or insteadof simpleshear.To
shelfof SW Greenland[ChianandLouden,1992]andthe depth constructa completerifting scenario,we require seismicdata
CHIAN AND LOUDEN:
CONTINENT-OCEAN
PSP PartiallySerpentinized
Peridotites
P
MG
8 •
CRUSTAL TRANSITION
9131
'
,,
Peridotites
Mafic Gneisses
,
,,'
'/
FG Felsic
Gneisses ,,'
SWG SW Greenland
0%
,
! ,/,
•
,
,
Lower Crust
/
,'
,
, ,'iLl
,'
0
[! •-..
Zone II
,
,
..
,
,
ß
ß
,
,
,
,
,
ß
,
ß
,
ß
ß
,
,
Zo'eIII
30,%
,
,
,
ß
,
,
Seismic Data'
Cont. U. Mantle
Cont. L. Crust
Lab Data:
6
¸
PyroxeneGabbro
r]Olivine
Gabbro
Serpentinite
3
4
Vs
5
(kin/s)
Figure 13. Compressional
wave velocity versusshear wave velocity showingfields for peridotires(P);
partiallyserpentinized
peridotires(PSP) with 0 and30 percentserpentine
pointsmarked;felsic gneisses(FG);
mafic gneisses(MG); reportedcrustalseismicdatafor continentallower crest(solidcircles)anduppermanfie
(vertical rectangles);and serpentinitedata at 0.6 kbar from ophiolitesamples(solid diamond)[Fountain and
Christensen,1989;Christensen,1966, 1978, 1989]. Field for SWG representsthe continentallower crustof
the SW Greenlandmargin [Chian and Louden,1992]. Also shownare laboratorydata for pyroxenegabbro
(opencircles)and olivine gabbro(openrectangles)[Spudichand Orcutt, 1980a;Christensenand Srnewing,
1981]. Fieldsfor zonesII andm are determinedfrom the lowercrustalseismicvelocitiesanduncertainties
(this
paper).
CHIAN
9132
AND LOUD•:
CONTINENT-OCEAN
on the conjugateLabrador margin. These data presenfiyare
beinganalyzedandwill be presentedin a subsequent
paper.
CRUSTAL
TRANSITION
has a very low velocity of 4.0-5.0 km/s and possibly
representsblock-faulteduppercontinentalcrust.
3. Thelowercrustin zone1TIhasV_ of-7.2 km/sandVs of
-4.1km/s.Such
velocities
possibly
extend
seaward
tothe
Conclusions
extinct ridge axis, and correspondto magnetic anomalies
Using wide-angle data and coincident normal-incidence youngerthan citron28. Seafloorspreadinglikely did not start
seismic reflection data, we have obtained a detailed two-
until chron 28.
dimensionalvelocity model for the completecrustaltransition
of the sediment-starved,SW Greenland continentalmargin
(Figure 10a). When convertedto densityvalues,this seismic
model offers a reasonablefit to the gravity data (Figure 12).
Interpretationof upper crustal smacturesis also compatible
with the change in magnetic anomalies and character of
4. A simpleshearmodelwithoutmagmatismis favoredfor the
initial rifting of the LabradorSea.The SW Greenlandmargin
possiblyacts as an upperplate margin,while the lower plate
lies on the conjugateLabradormargin.Seismicdataacrossthe
conjugatemargin of Labrador are needed to further test the
rifting model.
basement
structures.
The SW Greenlandcontinentalmargin is characterizedby
three distinct
zones:
1. In zone I, the 30-kin preexisting continental crust is
thinnedabruptlyto a thicknessof -3 kin. This thinnedcrustis
faulted and extends seaward to intersect a basementhigh,
interpretedas an elevatedserpentinizedperidotitc.
2.
Appendix: Error Analysisof the SeismicVelocity
Model
In the continent-ocean transition (zone II), the 4-5 km
lowercrust
extends
for50-80kmandexhibits
a bulkVp_
of
7.5+0.1km/sandVs of 4.0:L-0.2
km/s.Thisis interpreted
as
undercrusted
serpentinites.
The faulteduppercrustfor thiszone
In this sectionwe investigatethe errors of velocity and
depthobtainedfrom the two-dimensionalseismicmodeling.
We proceedby choosinga singlelayer, assming that all the
model parametersabove this layer are accurate.We then
perturbthe depthandvelocityfor the top boundaryof the layer
by dZo anddVo andperturbthedepthandvelocitygradient
for
the bottomboundaryby dZ 1 anddg. It is clearthat the
0.5
1.4
1.4
c-
"'
0
-0.5
(D
-1
-0.2
-0.1
0
0.1
Velocity Perturbation (km/s)
Figure
A1X2 contour
plotforperturbations
ofvelocity
and
depth
fortheseismic
phase
P3west
ofOBS
C
(Figure6). The contoursarenormalizedto the travel•ne endurance
(maximumerrorsin observed
traveltime)
of40ms.Thearea
enclosed
by X2=ldelimits
error
bounds
forthevelocity
anddepth
forthelower
crust.
0.2
CHIAN AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL TRANSITION
//
I
I
I
s•ue•u!pesJoj sNdAp
I
9133
9134
CHIAN
AND LOUDEN:
CONTINENT-OCEAN
CRUSTAL
TRANSITION
refracted
phasemainlycontrols
dZo anddVo andthatthewide- yields
angle reflection from the bottom boundarymainly controls
dZ1 anddg. Almostall the seismictraveltimeinformation
about this layer is includedin thesetwo seismicphases.The
exact travel time error analysisfor the layer parametersis thus
Io1= (v}-
<
(
five dimensional
(dZo, dZ1, dVo, dg, andtheroot-mean-square
traveltime misfit,Trms). To simplifythings,we performed If Vp=7.5km/s,dVp= :k0.1km/s,
Vs--4.0
krn/s,
anddVs= :k
theerroranalysis
for onephaseat a time,thatis, a (dZo, dVo 0.2 Icrn/s,the above-formulagives c•= 0.30 + 0.03.
andTrms)contour
plotfor therefracted
(diving)phasewithin
thelayer(omittingthelesssensitive
parameter
dg)anda (dZ1, Acknowledgments. This work was supportedby grantsfrom the
dg andTrms) contourplot for wide-angle
reflections
fromthe National Science and Engineering Research Council and the
bottomof the layer (assumingVo is accurateand therefore Departmentof Energy,Mines, and Resources,Canada.D. Chian was
partially supportedby a Graduate Fellowship from Dalhousie
aVo=0).
To quantify
theerrorbounds
of model
parameters,
a Z2
misfitis usedinstead
of Trms, where
University.We would like to thank the officersand crew of the C.S.S.
Hudsonand BordenChapman,Guy Fehn, Martin Uyesugi,and Neil
Hamilton for helping to collect the seismic data. Much of the
instrumentation
Toi andTci are the observedandcalculated
traveltimesat
selected points (i=l, 2.... n) of a specific seismic phase
whosetravel time pick has an uncertaintyof Uo (assumedto be
constant
forall i). Forvalues
of Z2 < 1, thecomputed
and
used at sea was on loan from the Atlantic Geoscience
Centre,BedfordInstituteof Oceanography.
Datafromthe oceanbottom
seismometers
were digitizedat the BedfordInstituteof Oceanography
with the helpof StevePerry.We thankS. Srivastava
for supplying
the
SOhObuoy
F seismicdata and C. Keen for early accessto a deep
multichannelreflectionprofile. Discussions
with and commentson the
manuscriptfrom J. Osler, S. Srivastava,C. Keen, L Reid, S. H olbrook,
and an anonymousreviewerhave beenuseful.
observedarrivals agree within the observeduncertaintiesof
the travel times. This criteria may be used for defining error
boundsof velocity, depth,and gradients.The boundsmight be References
reducedfurtherwhen consistentamplitudepatternsare modeled Balkwill, H.R., LabradorMargin:Structuralstyleandevolution,Mem.
in addition to travel times and where multiple raypathsfrom
12, editedby C. Beaumont
andA.J.Tankard,pp. 17-43,Can.Soc.of
different receiversoverlap.
Petrol Geol.,Calgary, 1987.
In the implementationof the error analysisfor the refracted Boillot, G., J. Girardeau,and J. Kornprobst,Rifting of the Galicia
Margin:Crustalthinninganderaplacement
of mantlerockson the
phase,the perturbation
of Zo is setto be uniformalongthe
seafloor,Proc. OceanDrill. ProgramSci. Results,103, 741-756,
top boundaryso that for all iterationsthe boundaryrelief is
kept the same.The perturbation
of Vo is constant
throughout
1988.
by
the layer so that thereis no gradientchangeduringiterations. Boillot,G., J., G. Feraud,M. Recq,andJ. Girardeau,"Undercrusting"
serpentinite
beneathrifted margins:The examplesof the west
The errors of the gradient can be estimatedby the second
Galiciamargin(Spain),Nature,341,523-525,1989a.
approach
for wide-anglereflections
in whichg andZ 1 are Boillot,
G., D. Mougenot,J. Girardeau,and E.L. Winrefer,Rifting
varied.
Figure
A1shows
ag2contour
plotforthephase
P3 processes
on the west Galicia margin, Spain, in Extensional
which is refractedfrom the high-velocity,lower crust for the
Tectonics
andStratigraphy
of theNorthAtlanticMargins,Mere.46,
final model in Figure 10a.
editedby AJ. TankardandH.R. Balkwill,pp. 363-377, Am. Assoc.
A similar approachcan be used for the error analysisof S
wavessuchas the PSP phasewhich is doublyconvertedat the
sediment/basementinterface. For PSS phases,with only a
single P-S conversion, the method for error analysis is
different
because
anextraunknown
Vsed, theS velocity
in the
sediments,is introduced.In this instance,we replacethe less
sensitivevariable, the depth perturbationdZo, with a
perturbation
of Vs inthesediment.
Figure
A2shows
a X2
contour
plotofd(Vp[Vs)
forthelower
crust
andd(Vp/Vs)
for
thetopmost
sediments.
Hereweusetheratiod(Vp/Vs)
instead
of Vs aloneformodeliterations
in ordertoreu•mthemodel
variations obtained from the generally better resolvedP
waves.
However,
thismethod
results
inm'mimum
Z2values
of
-4, indicating
thatstrictadherence
to theP wavemodelanda
uniform
Vi•IVs ratiowithineachlayerdoes
notresult
in
calculated-S
wave travel times that are able to fit the
observations within their estimated uncertainties. This
suggests
eitherthatourestimated
uncertainties
in theobserved
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undercrusted
serpentinite
beneath
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the layers,increasing
theuncertainty
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docan
beestimated.
D•ferenfiating
Vs
2
429-546, CRC Press,Boca Raton, Fla., 1989.
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CRUSTAL
TRANSITION
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B3I-I-4J1. (e-mail: [email protected],DepingChian)
(ReceivedFebruary3, 1993; revisedSeptember22, 1993;
acceptedNovember 30, 1993.)