A marine wide angle reflection/refraction profile adjacent Nova ...

Crustal Structure of the Meguma Terrane 

Offshore of Nova Scotia and its Tectonic 

Implications 

Version Dec 11/2006 

by 

H. R. Jackson 1 , D. Chian 1 , K. Louden 2 and M. Salisbury 1 

1 Geological Survey of Canada (Atlantic), Box 1006, Dartmouth, NS, Canada B2Y 4A2 

2 Department of Oceanography, Dalhousie University, Halifax, Nova Scotia, Canada B3H 4J1 

ESS contribution # 

E-mail: rujackso@nrcan.gc.ca; phone: 902-426-3791; fax: 902-426-6152

Abstract 

A 500 km long marine wide-angle reflection/ refraction (WAR) line was run parallel to the coast 

of Nova Scotia, on the Meguma Terrane. The terrane’s origin has been the subject of controversy 

exacerbated by the lack of exposure of lower crustal rocks. To provide constraints on the crust, P- and S- 

wave velocity models were developed using the data from 14 ocean bottom seismometers. Limited 

vertical incidence reflection profiles provide control for the modelling process. In situ laboratory 

measurements of the velocities and densities of rocks that are possible candidates for the crust were 

compared with the WAR velocities. The near seabed layer with velocity of 5.5-5.7 km/s and a thickness of 

3 km is interpreted as the Meguma Supergroup. A 10 km thick 4.3-5.0 km/s layer observed in the 

Orpheus Graben is inferred to be sedimentary strata. The velocity of the crust is 6.0-6.6 km/s, with a 

Poisson’s ratio of = 2.4 indicative of a felsic composition. Reflection profiles show a transparent upper 

crust that can be correlated with granitoid plutons. The lowermost crust is highly reflective. Both the 

Liscomb Complex and the Tangier granulite xenolith suite have high and low velocity members that could 

be the cause of the reflectivity. The subsurface contact between the Meguma and Avalon terranes is a 

southerly facing ramp with a length of about 60 km. The South Mountain Batholith of the Meguma 

Terrane is located to the southeast of the subsurface extent of the Avalon Terrane. Several hypotheses for 

the emplacement of the Meguma Terrane are examined. Delamination is one model that is consistent with 

the new constraints. 

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Introduction 

The Meguma Terrane (Fig. 1) consists principally of metasedimentary rocks of the Meguma 

Supergroup intruded by peraluminous granites. The basement to the extensive Cambro-Ordovician 

metasedimentary rocks (Fig. 2) is not exposed and its type, origin, and emplacement history have been 

difficult to determine (Greenough et al. 1999). The Meguma Terrane was the last to be added to the 

Appalachian Orogen, adjacent to the Avalon Terrane (Williams and Hatcher 1982). Keppie and 

Dallmeyer (1995) and Eberz et al. (1991) have suggested that the basement is Precambrian and of African 

origin or that the Meguma Terrane is thrust over the high-grade gneisses of the Avalon Terrane 

(Greenough et al. 1999). Murphy et al. (1999) hypothesized that hot spot activity was the cause of the 

voluminous granites and lesser mafic rocks. Pe-Piper and Jansa (1999) based on geochemical evidence 

advocate that the lower crust of the Meguma and Avalon terranes are similar and the voluminous igneous 

activity that produced the granites was produced by post-orogenic regional extensional. The available 

refraction data (Barrett et al. 1964) indicate a low overall crustal velocity incompatible with the Meguma 

Terrane overlying a high velocity basement due to Avalon crust or mafic underplating. 

The scientific objectives of this WAR experiment were 1) to develop a velocity model for the 

Meguma Terrane and 2) to determine the velocity of and depth relationship to the adjacent Avalon 

Terrane to gain insights into the enigmatic evolution of the Meguma Terrane. Bedrock samples, swath 

bathymetry and potential field data extend the surface geology offshore to aid in the interpretation of our 

WAR data. Limited seismic reflection profiles are available to compare the reflection characteristics of 

the crust to the WAR data. Significantly, our model velocities are evaluated against the velocities from 

3

laboratory measurements of candidate rocks for the crust. Gravity modelling is done to provide additional 

control on the velocity model. 

Geological and Geophysical Framework 

The Meguma Supergroup (Fig. 1) consists mostly of fine-grained siliciclastic turbidites and is of 

Gondwanan affinity based on its dispersal pattern (Schenk 1970, 1971). Its composite stratigraphic 

thickness exceeds 23 km but nowhere do all the units occur (Schenk 1995; 1997). It was deposited during 

the Cambrian (or older) to the Early Ordovician (Fig. 2). The Supergroup has a monotonous lithotectonic 

assemblage with similar structural style and level of exposure across the belt. There is no preservation of 

shelfal platform carbonate sequences and no exposure of deep-seated rocks. Deformation in the 

Supergroup was accommodated by chevron folding. 

Intruding it during from 380-370 Ma ( Keppie and Dallymeyer 1995) and representing at least 

25% of the exposed rocks are the peraluminous granites (Clark and Chatterjee 1992). The intrusions are 

massive; for instance, the South Mountain Batholith (SMB) (Fig. 1) is the largest in the Appalachian 

Orogen (Benn et al. 1997). Mafic intrusions are found only in the peripheral plutons and may contribute to 

the forming of the tonalite (Clarke et al. 2000) but are not a significant proportion of the outcrop. The 

mid-lower crust must have experienced high T/low P metamorphism to produce large volumes of 

granitoid rocks. It is unclear how to reconcile the history of the metasedimentary rocks with the extensive 

melting of the crust. 

Only three windows into the rocks of the deeper crust are available (Clarke at al. 1993): the 

granitoid intrusions; the Liscomb Complex consisting of granites, mafics, ortho-and paragneiss; and a 

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small group of lamprophyre dykes at Tangier, containing xenoliths, felsic and mafic granulites (Fig. 1). 

The composition of the granitoid intrusions indicates that the Meguma Supergroup cannot be the only 

source for the intrusions and there must be changes in rock type at depth (Clarke et al. 1988). 

Furthermore, the Sm/Nd systematics on the Devonian granites suggest they were formed from deep 

crustal material that is significantly younger than the Cambro-Ordovician metasedimentary rocks they 

intrude (Clarke et al. 1988). This age relationship reveals an inverted stratigraphy for the Meguma 

Terrane. 

The metamorphic grade of the Liscomb Complex is higher (amphibolite to granulite) than the 

Meguma Supergroup, but lesser than the lamprophyre dykes (Clarke et al. 1993), suggesting that it 

represents an intermediate level in the crust. Felsic xenoliths from the Tangier dyke (Fig. 1) have the 

appropriate composition to produce the granites. Their Tchur model ages are younger than the Meguma 

Supergroup flysch (Eberz et al. 1991) but consistent with the age of the granites. 

The Meguma Terrane was accreted to North America during the Acadian Orogeny in the Early 

Devonian (ca 400 Ma; Fig. 2) along the Cobequid-Chedabucto Fault (CCF) (Withjack et al. 1995, 

Webster et al. 1998, Pe-Piper and Piper, 2004). This fault can be traced from offshore Cape Breton across 

Nova Scotia into the Bay of Fundy. The CCF is exposed onshore and mapped using various geological 

and potential field data. This combined data set indicates that from surface to 6 km depth the fault is steep 

(Webster et al. 1998). Offshore of Cape Breton (Fig. 1) at the end of line 99-1 (Fig. 3), seismic reflection 

profile 89-1 (Marillier et al. 1989) shows the Avalon-Meguma contact as southward dipping. 

To place spatial limits on the Meguma Terrane the potential field and seismic information are 

examined. The regional gravity (Dehler and Roest 1998; Fig. 3) and magnetic anomaly maps (not shown; 

5

see Oakey and Dehler 1998) cover both the onshore and offshore. Onshore the potential field information 

is readily compared with the surface geology. It is extrapolated offshore by swath bathymetry, samples of 

basement from petroleum wells and short hard rock cores. Several regional negative gravity anomalies are 

observed in proximity to WAR 99-1 profile. The two largest are of similar shape and extent, one is 

centred on Nova Scotia and continues to the southern part of line 1 and the other is observed entirely 

offshore. The first anomaly coincides with the SMB and its near shore extent (Fig. 3) based on swath 

bathymetric mapping (Loncarevic et al. 1994). The major offshore gravity low is correlated with granitic 

rocks based on samples from D-26 and E-07 wells. A linear negative anomaly south of Cape Breton 

Island is associated with the thick accumulation of sedimentary rocks in the Orpheus Graben, the offshore 

extension of the CCF. 

Onshore the distribution of the Meguma Supergroup correlates with linear magnetic anomalies. 

This is due to mineralization along axial folds that occur evenly at about 8 km spacing. In contrast to the 

clastic rocks, the South Mountain and Musquodoboit batholiths that have similar mineralogy, petrology 

and geochemistry (Ham 1998) produce broad negative anomalies. Offshore the magnetic pattern 

associated with the Meguma Supergroup can be followed to the location of line 1 and beyond. This is 

substantiated by multibeam soundings that show linear bed forms associated with the Meguma 

Supergroup (Loncarevic et al. 1994). In addition, scientific drilling recovered a quartzite petrologically 

typical of the Meguma Supergroup (Pe-Piper and Loncarevic 1989) near the southwestern section of line 

99-1 (Fig. 3 Core 22). Further offshore the extent of the Supergroup is confirmed by phyllites and 

metaquartzites drilled in the offshore wells I-94, N-30, I-22, F-38 and E-53 (Fig. 3) (Jansa 1991; Pe-Piper 

and Jansa 1999). 

6

Three deep multichannel seismic reflection lines sampled on the Meguma Terrane on the margin 

(Keen et al. 1991a, b; Keen and Potter 1995). The profiles all showed a depth to the reflection Moho of 

between 10-12 s that is less than the adjacent Avalon Terrane at 12-14 s. Profile 86-5B (Marillier et al. 

1989) reveals the contact of the thicker Avalon crustal block with the Meguma Terrane south of the 

Orpheus Graben. Southward dipping reflectors cut the crust to the Moho separating the two crustal 

blocks. The cross-section of the crust on the Scotian shelf is displayed on deep multichannel reflection 

profile 88-1 (Keen et al. 1991b) and its extension 89-2 (Fig. 4). A transparent upper crust is observed to a 

depth of 15 km that is superseded by a moderately reflective crust to a depth of 25-30 km. The lower crust 

is most reflective from 33-38 km depth. 

None of the reflections can be correlated with the previous refraction results of Barrett et al. 

(1964). They measured velocities of 5.5, 6.10 and 8.11 km/s for the metasedimentary layer, crust and 

upper mantle. The small number of recording stations 4 and the shot spacing of 15 km suggested that the 

improvements in technology and interpretation techniques in the intervening 40 years would produce a 

higher resolution velocity depth profile if the line was rerun. 

Experiment 

On Hudson cruise 99-007, 14 digital ocean bottom seismometers (OBS) were launched and 

recovered (Fig. 3). The airgun array fired 3983 shots over a range of 503 km. The data quality was 

reasonable, except on the noisy OBS 13 deployed near the Orpheus Graben. OBS 1-8 recorded only about 

75% of the data due to a computer-programming fault. A second line (99-2) was positioned nearly 

perpendicular to the first (Jackson et al. 2000). 

7

The source for both lines was an array of 6 airguns with a total capacity of 6x16.4 L (6 x 1000 in 3 ) 

deployed at 20 m below sea surface and fired every 60 s at 4 knots for a shot interval of ~110 m. Global 

Positioning System was used to determine the shot positions. The untuned array produces a reverberatory 

signal; however, this signal has the high energy at 2-10 Hz needed for long distance propagation and 

develops shear waves. 

Shot times, corrected for clock drift in both the OBS and the firing system for the airguns, were 

combined with navigation to generate a shot table for each OBS. This in turn was used to convert the raw 

OBS data to SEGY format at sea. We process the SEGY data with a causal band-pass filter of 3-10 Hz. 

The data are displayed with gain increasing linearly with shot-receiver offset. When data quality are not 

the best, we apply coherency mixing to the data similar to that described by Chian and Louden (1992): an 

average of ~7 adjacent traces along a coherency-determined optimal slope between 2 and 8 km/s, which 

results in minimal distortion of wavelet. Due to the large number of OBS profiles to be presented, the 

WAR data are picked from the hydrophone (and occasionally geophone) channels, the height of which 

represent the estimated travel-time uncertainty, overlain with computed travel-time curves. Unfiltered data 

were sometimes used during digitization of the water waves and refractions from shallow sediment given 

their higher frequency content relative to deeper phases. For weak phases, such as some Moho reflections 

and refractions from lower crust and upper mantle, coherency mixing is found useful to properly identify 

arrival times and digitize them. 

Modelling 

The two-dimensional ray-tracing algorithm of Zelt and Smith (1992) was used to interpret the 

data. The model was constructed with 13 boundaries, each of which contained up to 38 boundary nodes to 

8

cover a distance of 530 km. These boundaries form 12 layers, each specified by up to 14 velocity nodes 

for the top and bottom of the layer. Model parameters are linearly interpolated between adjacent nodes 

during ray tracing and display. All OBSs were deployed along a line, except OBS 13. It is slightly off line 

and in the subsurface adjacent to a salt dome in the Orpheus Graben. OBS 13 recorded the worst quality 

data. 

We use “a, b, c and d” to label crustal boundaries, and Pa, Pb, Pc, and Pd for P-wave refractions 

from underneath them. Similarly, Sa, Sb, Sc, and Sd represent S-wave refractions. PmP and SmS 

represent P- and S-wave wide-angle reflections from the Moho. The vertical resolution general decreases 

with depth as the wavelengths increase (0.2-1.0 km). The lateral resolution in the upper crust in this study 

is determined primarily by the spacing of the OBS. The scale over which lateral variations are averaged in 

the upper crust is about half the receiver spacing or ~20 km (White 1989). Below the upper crust the 

resolution is limited by the ray path geometries associated with the arrivals modelled from these layers 

(Zelt and Forsyth 1994). Their generic results for the middle and lower crust are that velocities are not 

well constrained because most of the ray paths are reflected phases. In this study the middle crustal 

velocity of 6.4 km/s is constrained by refracted arrivals. However, in the lower crust and upper mantle, 

poor lateral resolution is caused by the refracted waves restricted to near-horizontal ray paths. An error 

analyses is performed, by perturbing model parameters until the computed arrival times no longer provide 

an adequate fit to the data on one or more OBS sections. In general, the error bounds from several OBS 

are smaller than or equal to that for a single OBS section. For example, the velocity uncertainty is 

sometimes found to be 0.15 km/s for a single OBS, but when several OBS are considered, it reduces to 

0.10 km/s. 

9

Sedimentary Layers 

Along profile 99-1, the lowest velocity layer is associated with the 70-km-wide,

instrument. On OBS 6 and 7, the velocity of the first arrivals is greater than 6.0 km/s. Eastward, modelling 

of the near offset arrivals on OBS 8, 9, 10 and 11 indicate that the velocities range from 5.9-6.1 km/s, 

below a velocity of about 5.5 km/s that is only 0.5 km thick. 

Crust and Moho 

For the upper crust, the observed amplitudes are usually strong on each OBS, modelled by a 

velocity of 6.0 km/s that increases downward to produce diving waves in the upper 5-km-thick crust (Fig. 

5). The amplitudes decrease dramatically beyond 50 km offset. This is modelled with a velocity of 6.0-6.3 

km/s and a low gradient in the upper 20 km depth. As an example, OBS 1 (Fig. 5) records clear P- and S- 

waves at

this layer the velocity structure is not directly constrained by our data, although OBS 12 and 14 do record 

deeper crustal arrivals as diving waves from underneath the graben. 

The crustal layers below “b” and “c” have a low vertical velocity gradient. All the data from OBS 

1-8 exhibit a relatively low velocity of 6.4-6.6 km/s between boundary “c” and the Moho. As a result, 

only upper crustal refractions and PmP can be clearly observed. In comparison, reflection profile 89-2 that 

crosses the refraction line near OBS 5 (Fig. 4) clearly shows a transparent upper crust, somewhat 

reflective middle crust, and more reflective lower crust. This highly reflective lower crust is also exhibited 

on many OBS in the Meguma Terrane (e.g. OBS 6 and 8; Figs. 6 and 7). 

The lack of turning rays in the Meguma Terrane crust makes the wide-angle reflection PmP 

clearly visible in nearly all our OBS profiles in this terrane (Fig. 8). The PmP reflection is not a single 

reflection event but occurs over an interval of at least a second. The lower crustal velocity is primarily 

determined based on the curvature of PmP and rare intra-crustal reflections. To determine the sensitivity 

of the seismic arrivals to lower crustal velocity, the velocities were perturbed by 0.2 km/s increments. 

The effect of this change is exemplified on OBS 6 (Meguma) and 10 (Avalon) shown in Figure 9. For 

OBS 6, the computed travel-time curve of PmP with the correct velocity of 6.6 km/s matches the 

curvature of observed amplitudes, whereas when this velocity is perturbed to 6.8 km/s, the computed 

curvature does not match. On the contrary, PmP on OBS 11 has to be modelled by a higher velocity of 

6.9 km/s. It can be seen that a velocity of 6.6 km/s does not generate the correct curvature (Fig. 9b). 

Therefore, the resolution of the lower crustal velocities is estimated to be better 0.2 km/s. This observed 

crustal velocity in Meguma Terrane is lower than in cratons but is consistent with the average Paleozoic 

crust (Holbrook et al. 1992). 

12

In the crust from the west end of the model to the Orpheus Graben, the velocity gradient is low. 

However reflections are recorded on most of the OBS (Fig .10) that indicate velocity discontinuities, 

labelled as “a”, “b”, “c” and “d” do occur. In the lower crust on OBS 9, 10 and 12 (Fig. 5) stronger 

reflections occur that are associated with a more significant velocity discontinuity. We also marked as 

solid segments on the model (Fig. 10) the range of the actual reflecting points from the Moho. 

OBS 14 (Fig. 11), located on the Avalon Terrane, recorded PmP and SmS, and Pn up to 350 km 

offset to the west. Modelling of the PmP of this OBS indicates a shallower (34 km depth) Moho 

underneath the Orpheus Graben. The Moho deepens westward to 38 km depth according to OBS 12, and 

shallows to 37 km depth further west. Pn is observed on OBS 6, 11 and 14, and its associated velocity 

from the uppermost mantle is estimated at 7.950.1 km/s. The picked arrivals for the complete data set 

overlain with computed travel time curves from the final model are shown in Figure 12. 

Shear waves 

Shear waves (or S-waves) refracted from upper and lower crust are observed on many OBS, such 

as OBS 1 (Fig. 5a, b), 8, and 14 (Fig. 11b). The picking of their arrivals is usually more difficult than for 

corresponding P-waves due to reverberations of the earlier P-waves. Iterative modelling suggests that we 

can obtain a general fit using a constant Vp/Vs ratio for each layer in the velocity model for Vp. Poisson’s 

ratios as derived from modelling the S-waves are shown in Figure 10 and 11. For the upper 10 to 15 km 

(laterally variable) the ratio is 0.21. It can be seen that, for the mid and lower crustal layers, the Poisson’s 

ratio is = 0.24. Barrett et al. (1964) determined values of Poisson’s ratio of 0.22 for entire crust of the 

Meguma Terrane consistent with ratios calculated here. 

13

Although this is a simplified description of the crust, computed travel-time curves of SmS, or the 

wide-angle S-wave reflection from the Moho, fit the observed data reasonably well. An example is OBS 

14 (Fig. 11b); the observed SmS is composed of many seconds of reverberations with its first arrival 

much weaker than later arrivals. After coherency mixing of the record section, the first arrival time of 

SmS can be picked near 145-170 km distance, while the general curvature of the SmS is clearly defined, 

and therefore modelled, from its multiple reverberations. 

Laboratory Velocity and Density Measurements 

To provide a quantitative basis for interpreting the seismic results discussed above in terms of 

petrology, the acoustic velocities of a large suite of rock samples from Nova Scotia that could be 

considered candidates for the middle and lower crust under the Meguma Terrane were measured in the 

laboratory at elevated confining pressures ranging from zero to 1 GPa to simulate in situ conditions. The 

samples included granulite facies rocks from the lower crustal Tangier xenolith suite (Giles and 

Chatterjee, 1987), high pressure rocks from the Liscomb Complex (Clarke et al. 1993), granites and 

tonalities from the Musquodoboit batholith and coastal and offshore plutons, plus samples of the Tangier 

dykes lamprophyres and the Meguma formation itself. 

The laboratory measurements were made using the pulse transmission technique of Birch (1960) 

and Christensen (1985) on 2.54 cm diameter, 3 to 6 cm long minicores drilled from the samples and 

trimmed to right cylinders. After the densities of the samples were determined using the immersion 

method, the minicores were jacketed with thin copper foil to prevent the pressure medium from invading 

the samples at high pressures. One MHz lead zirconate piezoelectric transducers and backup electrodes 

were placed on the ends of the minicores and the entire assembly was jacketed in gum rubber tubing. The 

14

samples were then placed in a hydrostatic pressure vessel at the GSC/ Dalhousie University High Pressure 

Laboratory in Halifax, N.S., Canada and compressional wave velocities (Vp) were measured at selected 

pressure intervals by sending a high voltage spike to the transmitting transducer and calculating the 

velocity from the length of the minicore and the time of flight of the signal between transducers. Identical 

procedures were used to measure shear wave velocities (Vs) except that lead zirconate titanate transducers 

with matching polarities were used instead of lead zirconate transducers. The velocities shown in Table I 

were recorded as the pressure decreased and are considered accurate to 0.5% for Vp and 1.0% for Vs 

(Christensen and Shaw, 1970), while the densities are considered accurate to + 0.005 g/cm 3 . As noted by 

many investigators, the measured velocities increase rapidly with pressure to about 200 MPa due to the 

closure of microcracks, but at higher pressures, the velocities tend to increase slowly and linearly with 

pressure in response to the intrinsic elastic properties of the constituent minerals of the rocks themselves. 

As can be seen in Figure 13, in which Vp and Vs are plotted versus density at 600 MPa, the in situ 

pressure at a mid-crustal depth of about 20 km, the granites and tonalites cluster fairly tightly about a 

mean density of 2.67 g/cm 3 and mean compressional and shear wave velocities of 6.18 and 3.59 km/s, 

respectively, while the mafic lithologies display a mean density of 2.88 g/cm 3 and mean compressional 

and shear wave velocities of 6.72 and 3.77 km/s, respectively. Similarly, at 1 GPa, the pressure at the 

base of the crust, the granites and tonalities display average compressional and shear wave velocities of 

6.25 and 3.62 km/s, respectively, while the mafic rocks average 6.76 and 3.80 km/s. Interestingly, the 

high pressure Liscomb and Tangier xenolith assemblages both display a much wider range in velocity, 

with Vp ranging from about 6.8 to 8.0 km/s at 600 MPa, with the highest velocity samples being 

metapelites rich in sillimanite and sapphirine. 

15

Gravity model 

To model the gravity anomaly a simple two-dimensional algorithm was used. We extract a 

profile from a dense regional data grid (Fig. 3)( Dehler and Roest 1998), interpolated along the refraction 

profile at an interval of about 1 km. The final velocity model in Figure 10 is divided into rectangular 

blocks, each bordered by boundary or velocity nodes of the velocity model. Each block is assumed to 

have a constant density , calculated from the average velocity (v) of the block, using an empirical 

relationship: = -0.6997 + 2.2302 v -0.598 v 2 + 0.07036 v 3 -0.0028311 v 4 . A maximum density of 3.30 

10 3 kg m -3 is assigned to the upper mantle. 

Although gravity modelling does not generate unique results, it can provide a valuable check on 

our velocity model, especially the Moho topography. For example, the modelled depth of the shallower 

Moho underneath the Orpheus Graben is sensitive to the longer wavelength of the computed gravity, 

whereas structures closer to the seabed (e.g. inside the graben) influence the shorter wavelengths of the 

computed gravity curves. In the area near the Orpheus Graben, the seismic methods have the weakest 

control. Therefore, we combined both methods to constrain the model. It can be seen that the shape of the 

observed and calculated gravity anomalies are similar (Fig. 10), with the greatest mismatch in the Orpheus 

Graben. This mismatch may be attributed to the oblique crossing of the WAR line with the graben, and to 

the lack of detailed control inside the graben. Gravity modelling of the Orpheus Gravity anomaly by 

Loncarevic et al. (1967) indicated a feature with a depth of 10 km similar to that determined by the 

seismic experiment. 

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Discussion 

The P-wave velocity of 5.5-5.8 km/s observed from the southwest end of the line to the Orpheus 

Graben (Fig. 10) is attributed to the metasedimentary Meguma Supergroup. This is deduced from the 

character of the magnetic anomalies supported by offshore geological samples and based on comparison 

with laboratory velocity measurements (Table I). The S-wave velocity of 3.2 km/s is lower than the range 

of velocities measured on the Meguma Supergroup samples 3.52-3.8 km/s (Table I). This may be due to 

heterogeneities of the lithologies within the Supergroup and the small number of laboratory 

measurements. 

The velocity model (Fig. 10) shows a maximum thickness of 3 km of this unit at the southwest 

segment of the line. Our velocity and thickness are similar to the range of measurements of Barrett et al. 

(1964) on this section of the line. Gravity modelling of the SMB (Benn et al. 1999) suggests that it was 

emplaced at the base of the Meguma Supergroup limiting the Supergroup’s thickness to 7 km. It is 

probable that uplift and erosion has significantly thinned this unit. Thus, we interpret the Meguma 

Supergroup as a thin sheet and suggest it is now a roof pendant. 

Orpheus Graben on the Cobequid-Chedabucto Fault and Avalon Terrane 

The most striking features of the near surface arrivals along the WAR profile are the delays in the 

first arrivals due to the thick low velocity section of 2.5-2.8 km/s and 4.4-5.6 km/s in the Orpheus Graben 

attributed to sedimentary rocks (Fig. 5b). The WAR velocities are constrained by well velocities for the 

upper 3 km and the gravity modelling suggest that the graben is a maximum of 10 km thick. The crustal 

velocities beneath the graben and to the east show evidence of being higher throughout the crust (Fig. 10). 

Similar velocities for the lower crust are recorded for the Avalon Terrane off Newfoundland (Hall et al. 

17

1998). The Avalon Terrane is measured to be 40 km deep on WAR line 91/2 along the Atlantic coast of 

Newfoundland and about 12 s or greater based on reflection profile 84-2 northeast of Newfoundland and 

offshore of southern New Brunswick (Keen et al. 1991a). 

The surface expression of the CCF (Webster et al., 1998; Pe-Piper and Piper 2004) is plotted on 

Figure 3. Its sub-surface shape at the northeastern end our study is based on reflection line 86-5 (Fig. 10). 

This reflection profile shows an inclined plane, similar to that required by our refraction data but 

perpendicular to it. The fault shape, crustal thickness and velocities suggest that Avalon Terrane underlies 

the Meguma Terrane as far southwest as OBS 7. Along profile 99-2 another control point for the extent of 

the Avalon Terrane is available. In the Bay of Fundy delimitation is based on reflection profiles 88-4 and 

88-2. We prefer the interpretation of Keen et al. (1991a) with the Avalon zone terminating near the 

beginning of profile 88-2 (Fig. 3). There is a crustal ramp on the reflection profile 88-4 that surfaces in the 

vicinity of the CCF (Withjack et al, 1995, Pe-Piper and Piper 2004) and tapers southward. The depth to 

Moho is 12s beneath the Avalon Terrane. The crustal wedge thins and terminates to the southward on 88- 

2 where a highly reflective lower crust and flat Moho at 10s are interpreted as the Meguma Terrane. 

Based on these control points from seismic data and extrapolating between them, the subsurface 

extent of the Avalon Terrane beneath the Meguma is also plotted on the gravity map (Fig. 3). These points 

connect parallel to the CCF. It is striking that the deep feather-edge of the Avalon Terrane parallels the 

negative gravity anomaly associated with the SMB and its offshore extent and terminates prior to the 

SMB. Geochemical data support this interpretation. Major element analysis of D-26 (Fig. 1) (Pe-Piper and 

Jansa 1999) falls within the average of the SMB analysis. In contrast, the samples from the F-22 well have 

Nd and Pb isotopes compositions similar to the Avalon Terrane (Pe-Piper and Jansa 1999). 

18

Crustal layer Meguma Terrane 

From the southern end of the profile 99-1 to the Orpheus Graben along the Meguma Terrane, the 

P-wave velocity of the crust increases from 6.0 km/s near the surface to 6.6 km/s above the Moho. This 

velocity range is compared to the laboratory measurement (Table I, Fig. 13). Assuming average heat flow 

values similar to those in the eastern United States (Blackwell, 1971; Lachenbruch and Sass, 1977), T will 

be about 300 o C at 20 km depth and 600 o C at the base of the crust. Assuming further that dVp/dT = -.39 

and -.55 x 10 -3 km/s o C, respectively, for granitoid and mafic rocks (Christensen and Mooney, 1995) and 

dVs/dT = -.15 and -.21 x 10 -3 km/s o C for the same lithologies (after Ji et al, 2002), Vp will be ~6.06 and 

6.01 km/s and Vs~3.53 and 3.50 km/s for granitoid rocks in the middle and lower crust; and Vp will be 

~6.57 and 6.46 km/s for mafic rocks in the middle and lower crust, while Vs will be ~3.69 and 3.65 km/s. 

Poisson’s ratio, which is relatively insensitive to P and T, since Vp and Vs tend to rise and fall together, is 

~ 0.24 for the granitoids and ~0.27 for the mafics, suggesting that granitoids dominate the crust and that 

mafics are only significant near the base of the section, perhaps interlayered with granitic rocks to cause 

the observed reflectivity. Another possibility if the temperature is slightly elevated is that the higher P- 

wave and Poisson’s ratio values near the base of the section could be caused by a phase change from to 

quartz which occurs at about 650 o C (Barruol 1993). 

Seismic reflection profile 89-2 (Fig. 4) shows a transparent upper crust to a depth of about 15 km 

compatible with a thick batholith (Fig. 4). Gravity modelling of the SMB indicates a thickness of 6 km 

(Benn et al. 1999). If for the batholith we estimate an average thickness of 10 km, a length of 500 and a 

width of 100 km, then the volume of granitoid rocks is 500,000 km 3 . There must be equally voluminous 

ultramafic restites below the batholith (eg Wolf and Wyllie 1994). One possible explanation for the lack 

19

of mafic crust is that it foundered into the mantle. Another is that the seismic and petrologic Mohos are 

not the same. Mafic-ultramafic rocks in the lower crust could be difficult to distinguish from the 

peridotites of the mantle. Further consideration of the origin of the crust is discussed later in this section. 

Development of the crustal structure 

We consider various hypotheses that can explain the inverted stratigraphy of the Meguma Terrane 

(Eberz et al., 1991, Clark et al 1993). One way for the older metasedimentary rocks to overly a younger 

crust is a tectonic break. Seismic reflection line 88-1 exhibits a spectacular group of dipping reflectors that 

could be consistent with thrusting (Keen et al. 1991b). However, the depth of these reflections is too great 

to be consistent with the 3 km thick section of metasedimentary rocks overlying the younger granitic 

crust. Furthermore, a tectonic thrust does not explain the even level exposure of the crust, and the lack of 

exposure of lower crustal rocks and ophiolites. 

A hot spot model has also been used to describe the voluminous short-lived magmatism of the 

northern Appalachians (Murphy et al. 1999). This model is not in agreement with the felsic nature of the 

crust. Furthermore, a hot spot would be expected to have underplated the crust with a mafic layer with 

velocities in the range of 7.0-7.6 km/s. We found no evidence of such a layer. 

Another explanation for the voluminous granitic rocks is regional extension (Pe-Piper and Jansa 

1999). The timing of the granitoid and mafic intrusions is late to post orogenic. The plutonism continued 

into the Carboniferous on the Scotian margin when the area was undergoing regional extension. Other 

evidence for extension is the wide across strike extent of the plutons, the ubiquitous extension in the 

Magdalen Basin, and the geochemical and isotopic variation of the granites that are attributed to 

20

decompression melting. Indeed, the elongated shape of the SMB is parallel to the CCF. However the other 

granite intrusives were emplaced along the northeast trending folds of the Meguma Supergroup (eg 

Keppie and Dallmeyer 1995) perpendicular to the CCF arguing against across strike extension to solely 

account for the emplacement of the granitoid bodies. Other evidence that extension does not account 

entirely for the formation of the granites are folding and shearing foliations in two batholiths that recorded 

tectonic deformation as they crystallized, indicating syn-tectonic emplacement (Benn et al. 1997). 

A more inclusive premise for the development of the Meguma Terrane that includes the 

generation of the granitoid rocks, the young source ages for the lower crust and the uplift history is 

detachment of the lower crust and upper lithosphere. Delamination (Nelson 1992) is the rapid mechanical 

thinning of the mantle beneath an orogenic belt (Fig. 14). Delamination has been previously proposed for 

the Meguma Terrane (Keppie and Dallmeyer 1995) and in the Gander zone in Newfoundland (Schofield 

and D’Lemos 2000). The sequence of events begins with a collision that thickens the crust and 

lithosphere, followed by detachment of the lower crust and lithosphere that initiates uplift and erosion 

followed by extension and subsidence. 

The Meguma Supergroup’s sedimentology is consistent with its formation in a volcanic arc 

setting. There are locally occurring volcaniclastic rocks and bulk chemical analysis suggests a strong 

volcanic component. In addition, the overlying Annapolis Supergroup is formed of subaerial 

volcaniclastic rocks (Schenk 1997). The meta-igneous xenoliths of Tangier dyke major and trace element 

compositions and ratios (Eberz et al. 1991), and the composition matrix (Giles and Chatterjee 1987) are 

consistent with crustal thickening in a continental volcanic arc. Greenough et al. (1989) based on U-Pb 

zircon and monazite dates for the xenoliths suggest that shortening between the Meguma and Avalon 

terranes occurred ~ 400 Ma and emplacement of the xenolith bearing dykes at ~370 Ma. 

21

The granulite-facies xenoliths found in the Tangier dyke are a feature diagnostic of delaminated 

terranes (Costa and Rey 1995). The P/ T conditions of these xenoliths suggest they come from greater 

depths than the present Moho (Giles and Chatterjee 1987, pressures of 12-14 kbars and temperatures of 

>1000 0 C). Based on the metamorphic grade of the xenoliths, Keppie and Dallmeyer (1995) advocate that 

the crust was 65 km thick. The subsequent foundering of the lower crust would explain the discrepancy 

with the present Moho depth. 

Delamination of the entire crustal root could result in complete melting of the lower crust (Nelson 

1992). Any region that has extensive granites such as observed in the Meguma Terrane may be inferred to 

have had temperatures that are hot enough for the lower crust to flow. Keppie and Dallymeyer (1995) 

propose that the emplacement of the granites about 30 Ma after initial convergence is consistent with the 

slow rise of the geotherms as the crust and lithosphere were thickened. They also hypothesize that the 

oldest pluton at 380-378 Ma was the time of the detachment of the lower crust. The composition of the 

metasedimentary xenoliths are consistent with the SMB being derived from them (Eberz et al. 1991). 

Seismic signatures of delamination are a crustal thickness that is thinner than the adjacent craton 

and a highly reflective lower crust (Nelson 1992) (Fig. 10 and 14). The consequence of detachment of the 

crustal root for the overlying rocks is uplift that is observed in the Meguma Terrane. Geobarometric data 

indicate that the present surface of the batholiths were emplaced at 6-10 km depth (Raeside and Mahoney 

1997). Based on a regional unconformity (Martel et al. 1993), within 13-23 million years of their 

emplacement the granitoid plutons were exhumed. The rate of erosion is consistent with ductile thinning 

of the crust and post-orogenic collapse. Exhumation erosion rates are < 5 km Ma -1 , in contrast to the faster 

exhumation due to synorogenic erosion and normal faulting. Following uplift a thermal sag basin is 

created in this model. This is compatible with the once wide spread distribution of 4-5 km Late 

22

Carboniferous to Early Permian strata over the Meguma Terrane (Ryan and Zentilli 1993). In summary, 

the delamination is one hypothesis that is consistent with many of the characteristic observed in the 

Meguma Terrane and explains its inverted stratigraphy. 

Conclusions 

A velocity layer of 5.5-5.7 km/s, attributed to the Meguma Supergroup, is 3 km thick or less 

adjacent the coast. The thinness of this section may be due to erosion. The velocity of the crust is 6.0-6.6 

km/s. Poisson’s ratio for the crust ranges from 0.21 to 0.24 indicates it is felsic. The upper portions of the 

crust are transparent on seismic reflection profiles. Granitoid rocks that form 25 % of the surface 

exposure on shore have both the appropriate P- and S -wave velocities and low reflectivity that make them 

a prime candidate for the upper crust. In the Meguma Terrane reflectivity is particularly high in the lower 

crust. From both the Liscomb Complex and the Tangier xenolith suite, there are high and low velocity 

rocks that could cause this. The suture between the Avalon and Meguma terranes is constrained to a 

southward ramp of 60 km in length. The SMB lies outboard and parallel to the edge of the subsurface 

extent of the Avalon Terrane. Four models for the tectonic history of the Meguma Terrane were 

considered. Components of several of the models are incorporated in a detachment hypothesis that is 

consistent with our velocity/depth model. 

Acknowledgments 

We would like to thank all the people who participated in the field experiment both on shore and 

at sea. We are particularly grateful to the Master, Captain Marsden, Officers and Crew of the CGGS 

Hudson whose excellent seamanship made launching and recovery of the airgun array and the ocean 

23

ottom seismometers possible in a variety of sea states. The Master and Crew of the forty foot CGGS 

Frank M Westin who launched the OBS on line 1 to save ship time for the larger vessel are thanked for 

their enthusiastic assistance. We thank Dr. S. Dehler who provided the plot files of the potential field data 

and B. Iuliucci who ran the samples through the high pressure laboratory. Dr. Clarke, Dr. S Dehler, and 

Dr. C. Hawkins provided comments that have substantially improved the manuscript. 

24

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33

Figures 

Figure 1) Location of wide-angle reflection /refraction profiles 99-1 and 2 relative to the onshore geology 

of Nova Scotia. The Cobequid-Chedabucto Fault (CCF) is shown separating the Meguma and Avalon 

terranes. SMB and MB in italics indicate the South Mountain and Musquodoboit batholiths, the sample 

locations at various rock types are indicated by: T for the Tangier dyke, L for the Liscomb Complex, M 

for the Musquodoboit Batholith, Me for the Meguma Supergroup. The dashed line 64-1 is the position of 

the refraction line of Barrett et al. (1964) and the thin solid line are deep multichannel reflection profiles. 

Figure 2) Geological time chart summarizing the evolution of the Meguma Terrane. 

Figure 3) Free Air (offshore) and Bouguer gravity (onshore) anomaly map. Hollow circles with labels 

indicate the location of wells and short bedrock cores. The crosses show the position of the Ocean Bottom 

Seismometers along the WAR profile. Deep multichannel seismic profiles in the region are shown as 

solid white lines. The solid black lines and dashed to dotted white lines represent respectively the surface 

and subsurface expression of the Cobequid-Chedabucto Fault. 

Figure 4) Multichannel seismic reflection profile 89-2 that crosses the WAR profiles 99-1 and 2. Strong 

coherency mixing was done to highlight the characteristics of the Meguma Terrane crust. Transparent 

upper crust (UC), reflective lower crust (LC) are labelled. The boundaries a, b, d and Moho are based on 

the WAR velocity model where it crosses the reflection profile. The seismic reflection profile was 

converted to depth using our velocity model. 

34

Figure 5) Computed travel time curves are overlain on the record sections of OBS 1and 12. The horizontal 

scale is shot-receiver distance (offset) and the vertical scale is the travel time using a reduction velocity of 

6.5 km/s for the P-wave. In 5b the P and S refracted waves from the metasedimentary layer and the upper 

crust above boundary a are shown. P and S wave velocities in km/s are shown as 5.5 (3.3). is Poison’s 

ratio followed by the P to S ratio in brackets. 5c OBS 12 at the edge of the Orpheus Graben shows the 

contrast in arrival patterns with OBS 1 located on the Meguma Terrane. 5d Refl LC indicates the zone of 

reflective lower crust. 

Figure 6) P-wave (a) and S-wave (c) record sections of OBS 6, showing typical features of the WAR 

Profiles along 99-1. The horizontal scale in the record section is shot-receiver distance (offset) and the 

vertical scale is the travel time using a reduction velocity of 6.5 km/s. The small vertical gradient below 

boundary a in (b) causes the Pa and Sa to terminate abruptly at ~ 60 k m offset. Abundant wide-angle 

reflections are produced between boundary d and the Moho modelled by lenses of high and low velocities. 

Section (b) illustrates the few arrivals between boundaries a and d. The numbers are velocities are in km/s. 

Computed travel time curves are overlain on the data in (a and c). 

Figure 7a) OBS 8 record section is overlain with the travel time curves. The horizontal scale in the record 

section is shot-receiver distance (offset) and the vertical scale is the travel time using a reduction velocity 

of 6.5 km/s. 7b is a plot of the ray tracing. 

Figure 8) Sections of OBS 6, 7, 9, 10 and 11 emphasizing the lower crust and Moho reflections. The large 

numbers at the upper edge of the record sections are the OBS numbers. All seismic profiles have the same 

35

horizontal and vertical scales while the velocity model is horizontally compressed. LC: lower crust. The 

dotted lines with arrows indicate the position of the record section on the profiles. 

Figure 9) The effects of different velocities in the lower crust on the computed travel time curves plotted 

on the record sections from OBS 6 and 11. The horizontal scale is shot-receiver distance (offset) and the 

vertical scale is the travel time using a reduction velocity of 6.5 km/s. The solid and dashed curves are 

computed PmP arrivals at various velocities. 

Figure 10) Calculated (solid line) and observed free air gravity anomalies are plotted across the top. The 

velocities are the largest numbers in km/s. Poisson’s ratio is shown as . The numbers in the ellipses 

indicate the particular OBS that constrain the layer at that point. The arrows on the top of the profile 

indicate where there is a reflection profile that crosses the WAR profile. The P-36 is the name of the well 

that is used to control the velocities in the Orpheus Graben The bottom section of the panel shows the 

seismic reflection profiles that are available at an angle to the WAR profiles to constrain the wide angle 

data. 

Figure 11) OBS 14 P- and S-wave arrivals (a) and (c) with the model for both shown in (b). 

Figure 12) Comparison of observed and calculated travel times for OBS 1 to 14, shown with the 

corresponding ray paths. The observed data are indicated by vertical bars with their heights representing 

the uncertainty of the picks; calculated data are indicated as solid lines. Vertical scale is in seconds with a 

reduction velocity of 6.5 km/s applied to the travel times and horizontal scale is offset in kilometers. The 

number of the OBS is indicated on the upper left corner of each section. OBS 13 was excluded because it 

was off line. 

36

Figure 13. Compressional (Vp) and Shear wave (Vs) velocity versus density for Nova Scotia samples at 

600 MPa. G, granitoid rocks; M, mafic rocks. Line L-T shows range of velocities for Liscomb and 

Tangier xenolith assemblages. Also shown are lines of constant acoustic impedance, Z; an impedance 

difference of 2.5 is sufficient to give a strong seismic reflection. 

Figure 14) A cartoon of the development of the Meguma Terrane in the context of a delamination model. 

The steps are: 1) the formation of the Meguma Terrane, 2) collison, 3) delamination, 4) extension and 5) 

collapse. 

37

TABLE I. LABORATORY VELOCITIES AND DENSITIES OF NOVA SCOTIA ROCK SUITES 

Sample Rock Type Density Mode Velocity (km/s) at Pressure (MPa) 

Meguma 

38 

g/cc 100 200 400 600 800 1000 

GO-1 greywacke 2.71 P 5.89 5.94 6.01 6.06 6.10 6.14 

2.71 S 3.71 3.74 3.77 3.78 3.79 3.80 

Hib-1 silicic greywacke 2.75 P 6.04 6.11 6.14 6.16 -- -- 

CH-1 staur-sill schist 2.84 P 6.01 6.09 6.16 6.20 -- -- 

HFX-3 slate 2.83 P 6.42 6.49 6.58 6.63 -- -- 

Musquodoboit batholith 

2.83 S 3.52 3.55 3.59 3.61 -- -- 

MI-1 granite 2.64 P 6.21 6.30 6.39 6.44 6.49 6.53 

Coastal and Offshore plutons 

2.64 S 3.66 3.71 3.75 3.76 3.77 3.77 

NS-17 granite 2.70 P 5.81 5.91 6.00 6.05 -- -- 

NS-20 granite 2.67 P 5.48 5.60 5.72 5.78 -- -- 

2.67 S 3.19 3.23 3.28 3.31 -- --

NS-21 granite 2.64 P 5.99 6.11 6.20 6.25 -- -- 

39 

2.64 S 3.53 3.59 3.62 3.65 -- -- 

NS-Y17 tonalite 2.74 P 5.99 6.11 6.20 6.25 -- -- 

2.74 S 3.36 3.43 3.51 3.56 -- -- 

NS-Y18 tonalite 2.66 P 5.67 5.88 6.00 6.08 -- -- 

NS-Y20A tonalite 2.72 P 6.00 6.11 6.22 6.25 -- -- 

2.72 S 3.55 3.59 3.64 3.67 -- -- 

NS-420B tonalite 2.71 P 6.36 6.48 6.61 6.69 -- -- 

Liscomb Complex 

LG-6 granite 2.62 P 5.79 5.94 6.06 6.12 6.16 6.17 

2.62 S 3.52 3.57 3.62 3.65 3.68 3.70 

LG-30 granite 2.63 P 5.91 6.02 6.10 6.14 6.18 6.21 

2.63 S 3.61 3.68 3.70 3.70 3.71 3.71 

LG-189 granite 2.63 P 5.99 6.03 6.15 6.23 6.28 6.30 

2.63 S 3.31 3.37 3.40 3.43 3.45 3.46 

LG-130-3 gneiss 2.84 P 5.97 6.10 6.21 6.25 6.27 6.31 

LG-98 augen gneiss 2.65 P 5.93 6.10 6.26 6.33 6.36 6.39 

2.65 S 3.60 3.65 3.69 3.71 3.73 3.75 

BI-98-1 gabbro 2.89 P 6.15 6.33 6.46 6.48 6.49 6.50 

2.89 S 3.27 3.32 3.37 3.38 3.39 3.39

LG-133-1 gneiss 2.82 P 6.27 6.36 6.45 6.50 6.52 6.54 

40 

2.82 S 3.66 3.70 3.75 3.77 3.78 3.78 

LG-178-1 gabbro 2.85 P 6.27 6.44 6.56 6.60 6.61 6.62 

2.85 S 3.52 3.60 3.68 3.71 3.72 3.72 

LG-172 gabbro-diorite 2.92 P 6.48 6.59 6.65 6.67 6.68 6.69 

2.92 S 3.67 3.72 3.76 3.78 3.79 3.80 

LG-146 gabbro 2.86 P 6.59 6.69 6.79 6.83 6.84 6.85 

2.86 S 3.76 3.78 3.80 3.81 3.83 3.84 

LG-165 gabbro 2.81 P 6.58 6.68 6.77 6.82 6.85 6.87 

2.81 S 3.63 3.68 3.72 3.74 3.76 3.77 

LG-162-2A mafic cumulate 2.93 P 6.67 6.84 7.00 7.10 7.15 7.19 

2.93 S 3.48 3.56 3.65 3.70 3.72 3.74 

LG-120 sill schist 2.97 P 7.53 7.60 7.68 7.74 7.81 7.87 

Tangier dyke lamprophyres 

2.97 S 4.13 4.18 4.20 4.23 4.26 4.28 

LG-197-1 lamprophyre 2.80 P 6.10 6.19 6.30 6.37 6.41 6.44 

2.80 S 3.47 3.52 3.57 3.60 3.61 3.61 

LG-194 lamprophyre 2.84 P 6.28 6.36 6.45 6.50 6.54 6.58 

Tangier dyke xenoliths 

2.84 S 3.61 3.64 3.66 3.67 3.69 3.70

LG-192x-23 metasediment 2.90 P 5.61 5.70 5.79 5.85 5.90 5.94 

sa+gt+sp>50% 2.90 S 3.36 3.46 3.53 3.57 3.59 3.59 


cpx+gt>50% 2.71 S 3.38 3.41 3.45 3.47 3.48 3.49 


opx+cpx>50% 2.91 S 3.31 3.38 3.47 3.52 3.55 3.57 


opx+gt>50% 2.82 S 3.59 3.65 3.69 3.71 3.72 3.73 


opx+gt>50% 2.82 S 3.57 3.58 3.60 3.62 3.63 3.64 

LG-196x-23 metabasalt 2.81 P 6.05 6.21 6.32 6.38 6.42 6.44 

amph>50% 2.81 S 3.68 3.73 3.79 3.83 3.85 3.86 

192x-6 metasediment 2.80 P 6.27 6.32 6.36 6.39 6.42 6.45 

sa+gt+sp>50% 2.80 S 3.41 3.48 3.55 3.59 3.62 3.65 


cpx+gt>50% 2.77 S 3.83 3.87 3.90 3.91 3.92 3.92 


cpx+gt>50% 2.88 S 3.80 3.81 3.83 3.84 3.85 3.85 

TD-98-1 metasediment 2.77 P 6.13 6.27 6.40 6.44 6.47 6.50 

cpx+gt>50% 2.77 S 3.69 3.73 3.75 3.75 3.76 3.76 

TD-92-3 metasediment 2.82 P 6.38 6.44 6.50 6.55 6.57 6.59 

cpx+gt>50% 2.82 S 3.63 3.65 3.68 3.70 3.71 3.71 

41

196x-23 metabasalt 2.79 P 6.19 6.41 6.52 6.56 6.60 6.63 

amph>50% 2.79 S 3.54 3.61 3.68 3.72 3.75 3.76 


amph>50% 2.83 S 3.49 3.56 3.64 3.66 3.69 3.69 

196x-26A metabasalt 2.80 P 6.28 6.46 6.58 6.64 6.68 6.71 

amph>50% 2.80 S 3.54 3.62 3.71 3.75 3.79 3.80 


opx+gt+sp>50% 2.86 S 3.47 3.49 3.50 3.52 3.53 3.54 


cpx>50% 3.01 S 3.81 3.87 3.92 3.96 3.98 3.99 

196x26 metabasalt 3.05 P 6.39 6.60 6.83 6.95 7.02 7.08 

cpx>50% 3.05 S 3.94 4.06 4.16 4.19 4.19 4.20 


sa+gt+sp>50% 3.00 S 3.81 3.84 3.88 3.90 3.91 3.92 


sa+gt+sp>50% 2.96 S 3.88 3.95 4.02 4.04 4.06 4.07 

abbreviations: silic, silicified; st, staurolite; and, andalusite; sill, sillimanite; gt, garnet; hb, hornblende; sa, sappharine; sp, 

spinel; opx, orthopyroxene; cpx, clinopyroxene; amph, amphibole; P, compressional wave; S, shear wave. 

42

88-2 

88-4 

CCF Bay of Fundy 

Carboniferous and Younger 

Sedimentary and Volcanic rocks 

Liscomb Complex 

Devonian Intrusive Rocks (Granites) 

Early Paleozoic Metasediments 

and Volcanics (Meguma Supergroup) 

North Mountain Basalt 

Rock Sampled for Vp-Vs 

Shelburne Dyke 

Recorders for 64-1 Line 

OBS Location for 99-1 and 99-2 

SMB 

MEGUMA AVALON 

64-1 

99-2 

Me 

be Gr 

CCF O pheus 

L 

M 

89-2 50 0 m 

PEI 

T 

64-2 

99-1 

1 2 3 4 6 7 8 9 10 11 

0 km 

MB 

99-2b 

100 

88-1 

Cp ae Breton 

r a n 

N ORTH AMERICA 12 

ATLANTIC 

OCEAN 

Fig. 1

TABLE OF FORMATIONS 

AGE 

Ma 

ERA PERIOD OROGENY 

66 

144 

208 

245 

286 

320 

360 

374 

387 

408 

438 

505 

570 

1000 

3000 

PROTEROZOIC PALEOZOIC MESOZOIC CZ 

QUATERNARY 

CRETACEOUS 

JURASSIC 

TRIASSIC 

PERMIAN 

CARBONI- 

FEROUS 

DEVONIAN 

SILURIAN 

ORDOVICIAN TACONIAN 

CAMBRIAN 

HADRYNIAN 

HELIKIAN 

HERCYNIAN 

deformation 

ACADIAN 

AVALONIAN 

GRENVILLIAN 

EVOLUTION OF MEGUMA TERRANE 

final rifting of Atlantic 

underplating SW Scotia Margin 

Stretching of Pangea; Shelburne dyke 

diabase; N. Mountain Basalt 

4 km erosion of Maritimes Basin 

subsidence 

formation of Maritimes Basin 

exhumation of granite 

intrusion of peraluminious granites, 

mafic dykes 

accretion of Meguma 

to N. America 

Meguma Group deposition 

formation of components of 

Meguma sediments: 

(Sm/Nd systematics and zircons) 

Fig. 2

67°W 65°W 63°W 61°W 59°W 57°W 

48°N 

88-2 

88-4 0 

km 

100 

Core 22 

99-1 B-93 

I-94 

99-2 

86-1 

PEI 

89-2 8 

1 6 7 

N-30 

E-7 

88-1 

CCCCFF 12 

P-36 

I-22 

865 - 

14 

D-26 

E-53 

F-38 

F-52 

50 30 10 -10 -30 -50 -70 

mGal 

89-1 

Fig. 3 

46°N 

44°N

Depth (km) 

0 

5 

10 

15 

20 

25 

30 

35 

40 

45 

SW NE 

MCS 89-2 

Refraction 99-1/99-2 

0 Distance (km) 20 

UC 

LC 

Moho 

a 

b 

d 

Fig. 4

PmP 

T-X/6.5 (s) 

12 

T-X/6.5(s) 

Depth (km) 

5 

4 

2 

1 

0 

0 

4 

W 

OBS 1 

(a) 

(d) 

(b) 

a 

0 km 10 

OBS 12 

(c) 

a 

b 

c 

d 

6.4 

6.3 

5.5(3.3) 

5.8(3.4) 

= .24(1.69) 

6.0(3.6) 

.05 

= .21(1.65) 

6.1(3.7) 

6.2(3.8) 

= .21(1.65) 

Distance (km) 

40 

Refl LC 

Moho 

6.5 

6.6 6.7-6.9 

6.9 8.0 

7.9 

0 50 100 150 200 250 300 

0 km 50 

Depth (km) -145-130 -115 -100 -85 -70 

Distance (km) 

OBS 2 

d 

a 

Sa 

Pa 

E 

Fig. 5

T-X/6.5(s) 

Depth(km) 

T-X/4.0(s) 

W E 

7 

6 

4 

2 

0 

40 

-150 

12 

10 

8 

6 

4 

2 

OBS 6 P-Wave 

(a) 

Pn 

(b) 

0 

-150 

a 

b 

c 

d 

-100 

OBS 6 S-Wave 

(c) 

-100 

6.2 

5.9 

6.5 

6.4 

-50 

-50 

6.3 

6.6 

PmP d 

SmS 

0 

50 

Distance (km) 

d 

a 

0 

50 

Distance (km) 

S2 

100 

100 

a 

6.7-6.9 

a 

150 

150 

Pn 

Fig. 6 

210 

Comlexity 

Near 

Orpheus 

Graben 

210

Depth (km) 

T-X/6.5 (s) 7 WE 

0 

OBS 8 d 

(a) 

d 

(b) 

a 

b 

c 

d 

Moho 

6.3 

6.4 

Refl LC 

6.5 

b 

6.6 

PmP 

40 

-100 -50 

0 50 100 150 

200 250 

Distance (km) 

a 

6.7 

6.9 

Fig. 7

Time-X/6.5 (s) 

Depth (km) 

6 

4 

PmP 

Refl 

LC 

2 

6 

4 

2 

W E 

6 

6 

4 

0 

10 

20 

30 

40 

10 

Avalon LC 

7 

2 

-120 -30 

a 

b 

c 

d 

Refl LC 

6.4 

6.3 

2 

-100 -40 

7 

m 

d 

Avalon LC 

7 

m 

2 

OBS 6 7 9 10 11 

6.5 

6.6 

? 

130 200 240 

0 90 

Avalon LC 

-150 -100 -50 0 50 100 150 200 250 300 350 

9 

9 

10 

7 

11 

0 50 100 20 60 100 120 

Distance (km) 

6.7 

6.9 

m2 ? 

Pn 

m 

6.7 

6.9 

Fig. 8

6 

4 

2 

W 6 

E 

Time - X/6.5 (sec) 

OBS 11 

5 

4 

3 

OBS 6 

Pn 

-100 -50 

Lower Crustal Velocity for Meguma Terrane: 

V=6.6 

6.8? 

20 50 100 120 

Distance (km) 

Lower Crustal Velocity for Avalon Terrane: 

V=6.9 

6.6? 

PmP 

PmP Fig. 9 

PmP (V= 6. 6?)

86-5b 89-2 99-1 2 

99- REFR99-1 

-100 mGal 100 

E 

MCS86-5b 

Avalon14 

4.2-4.8 

O. Graben 

P-36 

12 

11 

8 

7 

6 

MCS89-2 

REFR99-2 

5 

(Halifax) 

4 

5.5-5.8 

4.3 

5.9 

5.9 

6.5? 

5.5-5.8 

OBS 1 2 

.21(1.65) 

a 

1 

.21(1.65) 

W 

0 

.24 

1.69 

5 

Salts 

3 

6.0-6.1 .05 

6.2 .10 

6.1 

Meguma 

9 10 

.05 

10 

10 

6.2 

6.4 

6.3 

1 

60 . 

6. 

11 

9 

7 

8 

6 

6.2? 

4 

5 

0 

6 

5. 

9 

8 

.22 

 

b 

10 

6.7? 

6.8? 

Meguma Terrane 

2 

8 

.24(1.71) 

15 

10 

6.3 

6.4 

c 

20 

Marilliar et al. 

1994 

.24(1.71) 

Depth (km) 

2 

1 

9 

Avalon Terrane 

6.4 

8 

25 

5 

.1 

68 . 

8 

4 

6.5 

6 

10 

d 

30 

9 

14 

14 

.1 

10 

11 6.9 

7.95 .1 

6 11 14 

6 

5 

.1 

6.6 

3 4 

7 

6 

5 

.1 

.24(1.71) 

6.6 

35 

6.9? 

12 

9 

7 

2 

40 

0 50 100 150 200 250 300 350 

45 

-150 -100 -50 

N 

Avalon 

REFR 

99-1 

O. Graben 

Meguma 

REFR 

99-1/2 

SW NE S 

0 

0 

5 

UC 

10 

b 

10 

LC 

15 

20 

20 

25 

30 

30 

Depth (km) 

d 

MOHO 

40 

35 

40 

0 30 60 90 

45 

-150 -130 -110 -90 

MCS 86-5b 

MCS 89-2 

Fig. 10

T-X/7 (s) 

T-X/4.0(s) 

8 

6 

4 

2 

0 

45 

10 

8 

6 

4 

2 

0 

OBS 14 P-Wave 

(a) 

(b) 

a 

b 

c 

6.2(3.8) 

6.4 

=.21(1.65) 

d 6.5 

=.21(1.65) 

6.6 

OBS 14 S-Wave 

(c) 

SmS 

Pn 

8.0 

=.21(1.65) 

=.21(1. 65) 

63 . 

6.8 

0 50 100 150 200 250 300 350 

6.9 

d 

Distance (km) 

SmS d 

O Graben 

7.9 

PmP 

Fig .11

8 

6 

4 

2 

10 

20 

30 

40 

6 

4 

2 

10 

20 

30 

40 

8 

6 

4 

2 

10 

20 

30 

40 

6 

4 

2 

10 

20 

30 

40 

5 

3 

1 

-1 

10 

20 

30 

40 

1 

4 

7 

10 

14 

-100 0 100 

-100 0 100 200 

0 100 200 300 

2 

5 

-100 0 100 -100 0 100 

-100 0 100 

8 9 

11 

-100 0 100 200 -100 0 100 200 

3 

6 

12 

Fig .12

Velocity (km/sec) 

8 

7 

6 

5 

4 

3 

Vp 

Vs 

Lithology 

Meguma 

Musquodaboit, Plutons 

Liscomb granite 

Liscomb gneiss, schist 

Liscomb mafic 

Felsic Tangier xenoliths 

Mafic Tangier xenoliths 

Lamprophyre 

G 

G 

L-T 

M 

M 

Z= 17.5 

600MPa 

22.5 

2.5 3.0 

Density (g/cm3) 

3.5 

20 

Fig. 13

(1) 590-395 Ma 

(2) 395-388 Ma (3) ~ 355 Ma 

(4) 

SW 

Meguma deformation 

(5) 350-280 Ma 

35 km 

future 

Scotian margin 

Granites form 

385-370 Ma 

Meguma Supergroup Deposition 

CCF 

Meguma 

Avalon 

Meguma 

Meguma Avalon 

Carboniferous cover 

Meguma 

CCF CCF PP 

Avalon 

Delamination 

Moho 

150 km 

Exhumation of granites 

Moho 

40 km 

craton 

LEGEND 

CP – central pluton PP – peripheral pluton CCF– Cobequid Chedabucto Fault 

CP 

PP 

CCF 

Avalon 

Avalon 

NE 

Fig. 14

A marine wide angle reflection/refraction profile adjacent Nova ...

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