chapter 5 central atlantic magmatic province (camp): the
TRANSCRIPT
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CHAPTER 5
Central Atlantic Magmatic Province (CAMP): The Palisade Connection
Abstract
The 200 Ma Palisade Sills, exposed along the Hudson River in northeastern
North America are an expression of the Central Atlantic Magmatic Province (CAMP)
magmatism. On the basis of similar ages of eruption, Palisade Sill tholeiites have
been correlated to other CAMP exposures in four different continents. We provide an
isotopic tracer study of the Palisade Sill basalts and relate them to low-Ti (<2 wt %)
CAMP related tholeiites from North and South America, Europe, and West Africa.
We report Nd-Sr-Pb isotopic and multiple trace element data of nineteen basalts and
gabbros, three chilled margin basalts, and four sandstones spanning the entire length
and thickness of the Palisade Sill.
The Palisade Sill basalts of this study yield the typical composition of low-Ti
CAMP tholeiites with small LREE enrichments (LaN/SmN = 1.7 to 2.3), radiogenic Sr
and negative εNd(I) values (87Sr/87Sr(I) = 0.70668 to 0.71037; εNd(I) = -0.64 to -3.8), and
Pb-isotopic ratios (e.g. 206Pb/204Pb = 18.11 to 18.69) above the NHRL and subparallel
to it. The combined geochemical data of the Palisade Sill basalts and their correlation
with other low-Ti CAMP related lavas imply a slightly enriched mantle source with
likely contamination by the continental crust and with no contribution from an EM-I
or depleted MORB like component.
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Although both plume and lithospheric mantle sources have been previously
proposed for the low-Ti CAMP lavas, we use the geochemical data of our study to
propose the source of the Palisade Sill basalts as well as other low-Ti CAMP
tholeiites to be derived from ~15% melting of a slightly depleted spinel peridotite
which experienced up to ~20% contamination by the continental crust prior to or
during the emplacement of these lavas. We suggest the geodynamic emplacement
processes of the CAMP to be similar to the present day East African Rift System
which is sourced from the African superplume via multiple feeder stems that span a
large aerial extent.
5.1. Introduction
The extensively studied 200Ma Central Atlantic Magmatic Province (CAMP,
Fig. 5.1a) is considered to be the biggest Large Igneous Province (LIP) on this planet
covering up to 7 X 106 km2 (Marzoli et al., 1999; Olsen, 1999; Hames et al., 2000).
This igneous province has been linked to the early Mesozoic initial opening of the
Central Atlantic Ocean (Dalrymple et al., 1975; Bertrand et al., 1982; Alibert, 1985;
Dupuy et al., 1988; Bertrand, 1991; Sebai et al., 1991; Deckart et al., 1997; Hames et
al., 2000; McHone, 2000; Cebria et al., 2003; DeMin et al., 2003). The opening of
the Central Atlantic Ocean fragmented the CAMP into several segments that occur on
four different tectonic plates today. This magmatic event has been compared to
formation of flood basalt provinces such as the Siberian and Deccan Traps, in that
each may be genetically linked to a global faunal extinction (e.g. Olsen, 1999). For
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CAMP this is the event recorded at the Triassic-Jurassic boundary (Marzoli et al.,
1999). Although many authors have done extensive work on CAMP-related tectonics
and magmatic processes, there is still no consensus on its origin, and many aspects of
the CAMP remain controversial (e.g. Deckart et al., 1997; Olsen, 1999; Hames et al.,
2000; McHone, 2000; Puffer, 2002).
The CAMP related LIP is different from others in that it constitutes almost
entirely of dykes and Sills with scarce volcanic outflows (Cebria et al., 2003). In the
pre-Atlantic Ocean reconstruction, this dike swarm defines an overall radiating
pattern extending nearly 300 km from its focal point (May, 1971; Ernst et al., 1995).
Although the distribution of dikes on a local scale shows more complex patterns (e.g.
Bertrand, 1991; Hames et al., 2000), on a large scale it represents the best example on
this planet of a complete radiating dike swarm system (Greenough and Hodych, 1990;
Ernst et al., 1995; Dalziel et al., 2000; Beutel, 2009), comparable only with some of
the radial dike swarms of Venus (Ernst et al., 1995).
Extensive Ar-Ar dating and limited U-Pb dating have shown this even to be of
short duration with magmatism in all regions occurring within a few million years at
200 Ma (e.g. Dunning and Hodych, 1990; Sebai et al., 1991; Hodych and Dunning,
1992; Deckart et al., 1997; Olsen, 1997; Marzoli et al., 1999; Olsen, 1999; Hames et
al., 2000). Although non-plume models have been considered for the CAMP event
(McHone, 2000), a mantle plume appears to be necessary to explain the radiating dike
pattern and the generation of such a huge area of basaltic magmatism within only a
few million years (e.g. Wilson, 1997; Ernst and Buchan, 2002).
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Figure 5.1a. General distribution of Early Jurassic tholeiitic dikes, lavas and sills in a
pre-drift reconstruction at 200Ma (May, 1981; Schermerhorn et al., 1978; Belleni et
al., 1990; McHone, 2000). The square represents the Palisade Sill region.
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The study of regionally extensive dike and sill systems represents one of the
fundamental tools in the analyses of LIPs to clarify their origin (e.g. plume vs. non-
plume) and geodynamic setting. In this context the detailed geochemical study of the
Palisade Sill reported here may shed light on some of the unsolved questions linking
LIPs and regional dike swarms, such as (1) the association of radiating dikes with
either sublithospheric plume impingement (Ernst et al., 1995) or mantle insulation
beneath highly refractory cratons (Yale and Carpenter, 1998) and (2) vertical
emplacement versus far reaching (>3000 km) and nearly instantaneous (in less than a
few million years) lateral migration of magma from its source (Greenough and
Hodych, 1990; Ernst et al., 1995; Elliot et al., 1999).
It is generally agreed that the Palisade Sill basalts are an expression of the
CAMP magmatism that related to the earliest stages of the opening of the Central
Atlantic (Ernst et al., 1995; Oyarzun et al., 1997; Wilson, 1997). In terms of trace
elements and multiple element isotopic systematics, this poorly studied area of
CAMP, presently located in northeastern America is similar to the voluminous low-Ti
tholeiites associated with CAMP and distinctly different from the smaller volumes of
high-Ti magmatism related to the CAMP.
In this study we report the detailed trace element concentrations and Nd-Sr-
Pb-isotopic ratios of nineteen basalts and gabbros, three chilled margin basalts, and
four sandstones from the entire length of the Palisade Sill (Fig. 5.1b). These
geochemical data are essential to understand the relationship between mantle
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geodynamic processes involved in the generation of the CAMP tholeiites prior to the
formation of the of the Atlantic Ocean crust.
5.2. Geological History of the Palisade Sills
The initial fragmentation of the Pangaea supercontinent was accompanied by
extensive tholeiitic magmatism now represented by Sills, dikes and minor lava flows
in four continents along both sides of the Central Atlantic Ocean, on the eastern
margin of north America (between Nova Scotia and Florida), Western Europe (Spain
and France), West Africa (Morocco to Ivory Coast), and northern south America
(French Guyana, Surinam and Brazil) (e.g. Cebria et al., 2003).
The Palisades Sill is one of several diabase intrusions located in the system of
Eastern North American rift basins, formed by significant crustal extension during the
break-up of Pangea during the late Triassic. This early Jurassic Palisade Sill that
intruded into the Triassic continental sedimentary rocks of the Newark basin has been
cites as the classic example of a vertically differentiated Sill (e.g. Lewis, 1908;
Walker, 1940; 1956; Carmichael et al., 1974; Shirley, 1987; Husch, 1992; Gorring
and Naslund, 1995). The Palisade Sill and other sills in the Newark basin are used as
markers to define the Triassic-Jurassic boundary (Dunning and Hodych, 1990; Kent
et al., 1991). The intrusion is a ~200Ma (Erickson and Kulp, 1961; Dallmeyer, 1975;
Dunning and Hodych, 1990), 300m thick diabase Sill intruded into sandstones and
arkoses of the Newark basin (Walker, 1969b). The outcrop extends for 80 km in a
north-south direction from central Staten Island in New York, through a significant
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Figure 5.1b. Distribution of early Jurassic igneous rocks throughout the northern Newark
Basin (Puffer, et al., 2009). The location of the Palisade sill basalts discussed in this study
relative to eastern North America is shown in the inset.
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portion of New Jersey, along the western bank of the Hudson River (Fig. 5.1b)
(Naslund, 1998). Here it turns westward and is discordant with the local strata and
hence referred to as a dike locally (Walker, 1969b). It may connect under cover with
the Rocky Hill and Lambertville Sills to the south for a total strike length of 150 kms
(Husch, 1992). Field relations and petrography have been described by a number of
authors (Lewis, 1908; Walker, 1940; Walker, 1969a; Shirley, 1987; Husch, 1992;
Steiner et al., 1992).
Several authors have done major element analyses of the various layers of the
Palisade Sill to understand compaction and differentiation processes experienced by
this Sill after emplacement (e.g. Shirley, 1987; Husch, 1992; Puffer, 2002). However
very limited trace element and isotopic geochemistry related work has been done in
this region (Pegram, 1990).
5.3. Analytical Results
In this section we present the geochemical results of the Palisades Sill samples
of this study that include 17 basalts and gabbros, 4 sandstones, and 3 chilled margins.
These data comprise multiple trace element concentrations including the rare earths,
and the isotopic compositions of Nd, Sr, and Pb. The data are presented in tables 5.1-
5.2 and figures 5.2-5.7 and are compared with similar data obtained from literature on
volcanic rocks related to CAMP activity from North America, South America,
Europe, Africa, and Canada. Analytical methods are described in Appendix-2.
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Table 5.1. Trace element concentration of the basalts and sandstones from the
Palisade Sill analyzed for this study.
PS-01 PS-02 PS-03 PS-05 PS-07 PS-08 PS-10A PS-10B
Basalt Basalt Basalt Basalt (Dyke)
Basalt (Dyke) Gabbro Basalt
(Massive) Basalt
(Weathered) Rb 28.4 19.8 23.1 24.7 19.0 26.8 19.0 20.8 Ba 89 90 169 88 180 207 159 164 Sr 147 154 171 144 165 168 159 123 Pb 5.27 5.38 14.1 16.4 20.1 16.8 6.24 15.0 La 11.4 9.03 9.38 8.55 13.1 9.96 10.5 7.29 Ce 24.5 19.4 20.2 19.2 24.9 22.2 21.8 15.7 Pr 3.07 2.58 2.69 2.60 3.60 2.99 2.88 2.05 Nd 12.3 10.9 11.2 11.1 15.1 12.8 12.2 8.57 Eu 0.98 0.96 0.99 0.91 1.15 1.14 0.99 0.77 Sm 3.14 2.92 2.93 3.00 3.95 3.46 3.13 2.24 Gd 3.39 3.09 3.14 3.16 4.18 3.55 3.35 2.34 Tb 0.57 0.54 0.55 0.56 0.74 0.65 0.57 0.41 Dy 3.63 3.34 3.52 3.55 4.55 4.11 3.50 2.66 Ho 0.77 0.71 0.73 0.76 0.97 0.88 0.75 0.57 Er 2.10 1.93 1.94 2.01 2.56 2.38 2.02 1.52 Tm 0.32 0.29 0.30 0.30 0.38 0.36 0.30 0.24 Yb 1.99 1.84 1.92 1.91 2.41 2.30 1.97 1.51 Lu 0.28 0.27 0.28 0.28 0.36 0.34 0.29 0.22 Y 21.3 19.85 19.9 20.5 27.7 23.9 21.2 15.8 Th 1.85 1.79 2.10 1.74 1.97 2.26 2.23 1.68 U 0.64 0.43 0.51 0.43 0.61 0.52 0.51 0.45 Zr 85 89 99 74 95 103 95 84 Hf 2.27 2.36 2.74 2.02 2.65 2.75 2.47 2.11 Nb 9.75 11.3 15.1 10.2 12.2 14.0 9.19 7.45 Ta 0.77 0.72 0.95 0.66 0.96 0.89 0.58 0.91 Sc 36.3 33.2 31.5 32.1 39.6 32.8 33.2 22.7 V 311 313 321 325 351 303 332 252
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Table 5.1 continued
PS-11 PS-12 PS-13 PS-14 PS-17 PS-20 PS-21 PS-23 PS-24 Gabbro Basalt Basalt Basalt Gabbro Basalt Basalt Basalt Basalt
Rb 33.5 17.2 76 25.0 31.5 22.0 13.3 23.5 11.1 Ba 213 92 392 190 248 216 88 219 72 Sr 208 157 244 171 236 207 205 182 150 Pb 9.47 31.7 133 8.88 51 14.1 34.8 35.8 6.20 La 10.2 7.11 21.6 10.3 16.0 12.7 10.3 12.0 7.76 Ce 21.7 17.1 46.1 22.8 32.8 27.6 24.1 25.9 17.4 Pr 2.84 2.01 5.95 3.04 4.20 3.63 3.04 3.39 2.40 Nd 11.9 8.42 24.4 12.9 17.0 15.4 12.6 13.8 10.2 Eu 1.14 0.73 1.72 1.08 1.41 1.31 1.13 1.32 0.93 Sm 3.08 2.12 6.17 3.49 4.30 4.08 3.44 3.80 2.95 Gd 3.27 2.18 6.56 3.71 4.55 4.43 3.63 4.20 3.20 Tb 0.57 0.36 1.10 0.65 0.73 0.77 0.62 0.73 0.56 Dy 3.55 2.25 6.59 4.04 4.36 4.71 3.89 4.56 3.54 Ho 0.76 0.46 1.42 0.85 0.90 1.00 0.81 0.98 0.75 Er 1.99 1.24 3.76 2.26 2.48 2.66 2.17 2.64 2.01 Tm 0.31 0.18 0.58 0.35 0.36 0.40 0.33 0.41 0.31 Yb 1.97 1.11 3.66 2.20 2.29 2.59 2.09 2.63 1.96 Lu 0.28 0.16 0.53 0.33 0.32 0.38 0.30 0.38 0.28 Y 20.4 10.1 40.2 22.8 26.4 27.1 21.2 26.7 19.0 Th 2.08 1.34 5.51 2.14 3.48 2.67 2.07 2.13 1.44 U 0.50 0.52 1.31 0.51 1.06 0.61 0.48 0.54 0.34 Zr 94 106 209 101 164 126 83.1 105 76.2 Hf 2.57 2.85 5.08 2.81 3.99 3.26 2.25 2.79 2.19 Nb 8.71 10.6 15.8 9.76 12.7 11.2 8.07 9.71 6.19 Ta 0.54 0.65 1.07 0.63 0.81 0.68 0.51 0.58 0.38 Sc 32.9 10.8 22.4 33.9 20.9 32.4 25.0 39.4 34.8 V 338 278 274 345 249 373 247 368 316
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Table 5.1 continued
PS-24 PS-04 PS-15 PS-22 PS-06 PS-09 PS-16 PS-19
Basalt Chilled Margin
Chilled Margin
Chilled Margin
Silicified Sandstone
Sand- stone
Sand- stone
Sand- stone
Rb 11.1 187 123 25.1 93 76 69 39 Ba 72 558 301 261 885 652 940 262 Sr 150 134 123 205 289 86 331 94 Pb 6.20 39.8 20.1 14.2 18.5 25.5 75 21.3 La 7.76 31.0 18.4 11.7 33.6 55.6 16.9 6.33 Ce 17.4 82 33.3 25.2 70 218 34.1 18.4 Pr 2.40 8.04 3.96 3.26 8.12 13.7 4.80 3.17 Nd 10.2 30.5 13.9 13.7 29.2 49.6 18.2 16.1 Eu 0.93 1.39 0.94 1.24 1.36 1.25 1.07 0.82 Sm 2.95 6.86 2.70 14.3 6.11 9.92 4.14 6.19 Gd 3.20 6.79 2.56 3.94 5.63 8.61 3.99 6.00 Tb 0.56 0.99 0.34 0.67 0.80 1.01 0.63 1.22 Dy 3.54 5.44 1.93 4.11 4.50 4.42 3.75 7.80 Ho 0.75 1.02 0.40 0.90 0.91 0.71 0.78 1.62 Er 2.01 2.51 1.09 2.38 2.42 1.82 2.11 4.30 Tm 0.31 0.35 0.18 0.36 0.36 0.24 0.33 0.67 Yb 1.96 2.10 1.11 2.31 2.27 1.64 2.12 4.08 Lu 0.28 0.29 0.17 0.33 0.33 0.25 0.31 0.54 Y 19.0 35.0 11.5 24.0 25.4 18.5 23.2 44.3 Th 1.44 16.8 6.23 2.34 13.8 25.2 6.61 8.82 U 0.34 2.75 1.59 0.56 2.66 2.37 2.45 1.13 Zr 76.2 77 22.2 118 127 105 168 68 Hf 2.19 1.92 0.72 3.00 4.11 3.45 4.91 2.31 Nb 6.19 31.2 11.3 10.2 5.74 13.2 6.43 8.33 Ta 0.38 2.15 0.79 0.61 0.43 0.88 0.40 0.62 Sc 34.8 16.3 9.24 31.0 2.98 8.50 2.68 6.18 V 316 153 151 374 24.0 25.4 22.5 25.0
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5.3.1. Trace Element Geochemistry
Trace element data for all the Palisades Sill basalts and gabbros and related
sandstones and chilled margins of this study are presented in Table 5.1. The
Chondrite normalized (Evensen et al., 1978) rare earth element (REE) patterns for the
basalts and gabbros are shown in figure 5.2a, while those of the sandstones and
chilled margin basalts are shown in figure 5.2b. The basalts and gabbros show
uniform patterns with slight light rare earth element (LREE) enrichment (LaN/SmN =
1.7 to 2.3) and a relatively gentler slope for the heavy rare earth elements (HREE)
(GdN/YbN = 1.3 to 1.6). In contrast to the basalts and gabbros, sandstones and chilled
margins show a much wider range of chondrite normalized REE patterns, especially
LREEs (LaN/SmN = 1.1 to 23). These sandstones and chilled margin basalts show
LREE enrichment, ~20-100 times that of chondrite, HREE ~10-40 times that of
chondrite (GdN/YbN = 1.2 to 4.3), and strong negative Eu anomalies (Fig. 5.2b).
Twenty-two compatible and incompatible trace element concentration patterns
for the Palisades Sill lavas and associated sandstones and chilled margin basalts are
shown normalized to primitive mantle (Sun and McDonough, 1989) in figures 5.3a
and b, respectively. Primitive mantle normalized basalts and gabbros show low Ba,
high U, Pb, Zr-Hf concentrations, mildly negative Nb-Ta anomalies, and gently
sloping HREEs. In general, basalts and gabbros of the Palisades Sill are similar to and
~30 times more enriched than primitive mantle. Sandstones and chilled margin
basalts (Fig. 5.3b) have enrichments and depletions in the same elements as the
Palisade sill basalts and gabbros, but the magnitude of these enrichments and
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Figure 5.2. Chondrite normalized REE patterns of basalts from the Palisades sill
samples of this study: (a) Basalts and Gabbros and (b) sandstones (open circles) and
chilled margins (gray filled circles). Palisade Sill samples of this study have been
compared to high and low-Ti CAMP basalts from Europe (Alibert, 1985; Demant et
al., 1996), Brazil (DeMin et al., 2003), average Guinea, and average Guyana (Deckart
et al., 1997).
226
Figure 5.3. Multiple trace element concentrations normalized to primitive mantle for
the Palisades sill samples of this study: (a) Basalts and Gabbros and (b) sandstones
and chilled margins. Symbols as in figure 5.2
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depletions are much larger for the sandstones and chilled margin basalts. In addition,
they also have negative Sr and positive Th peaks. The Nb-Ta anomaly is much
stronger in the sandstones than in the chilled margin basalts.
The La/Ba (0.02-0.08) and La/Nb (0.8-8.3) ratios of the Palisades Sill data as
well as associated sandstones and chilled margins are plotted in figure 5.4a and
compared to various mantle reservoirs (Saunders et al., 1992) as well as CAMP data
from France (Jourdan et al., 2003), Guinea, and Guyana (Deckart et al., 1997). Notice
the close correspondence of the Palisades basalts and gabbros with OIBs, and low-Ti
CAMP related tholeiites from Guinea and France.
Various trace element concentrations and ratios have been plotted in figure 4.
The plot of Nb (ppm) versus Nb/Ta (Fig. 5.4b) compares the basalts and gabbros of
this study with sediments (Govindaraju, 1994; Vroon et al., 1995), ocean island
basalts (OIB) (Clague and Frey, 1982; Frey and Clague, 1983; Palacz and Saunders,
1986; Weaver et al., 1987; Chauvel and Hofmann, 1992), average mid-ocean ridge
basalt (MORB) and primitive mantle (PM) (Sun and McDonough, 1989). Also shown
is calculated variation of Nb with Nb/Ta in melt and residue at variable degrees of
batch melting of a peridotitic source with primitive mantle Nb (0.713 ppm) and Ta
(0.041 ppm) concentrations. Peridotite melting models are based on starting
composition of olivine (60%), orthopyroxene (25%), clinopyroxene, (10%), and
garnet (5%). ‘D’ values are from Green et al. (1989).
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Figure 5.4a. Comparison of the Palisade sill basalts and associated sandstones and
chilled margins with high and low-TI CAMP basalts from Guyana, Guinea (Deckart
et al., 1997), Brazil (DeMin et al., 2003), and France (Jourdan et al., 2003) in the
diagram La/Ba versus La/Nb (Saunders et al., 1992; Nomade et al., 2002). DM:
depleted mantle; CC: continental crust; PM: primitive mantle; OIB: ocean island
basalts.
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Figure 5.4b. Plot of Nb/Ta versus Nb (ppm) of the Palisade sill basalts and gabbros
of this study compared with sediments (Govindaraju, 1994, Vroon et al 1995), ocean
island basalts (OIB) (Clague and Frey, 1982; Frey and Clague, 1983; Palacz and
Saunders, 1986; Weaver et al., 1987; Chauvel et al., 1992), average mid-ocean ridge
basalt (MORB) and primitive mantle (PM) (Sun and McDonough, 1989). Also shown
is calculated variation of Nb and Nb/Ta in melt and residue at variable degrees of
melting of a peridotitic source (see text for details).
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5.3.2. Nd-Sr-Pb Geochemistry
Rb-Sr and Sm-Nd isotope systematics data for all the rocks of this study are
reported in Table 5.2. The initial εNd and 87Sr/86Sr(I) ratios of all the Palisade Sill
basalts of this study are plotted in figure 5.5. Initial εNd and 87Sr/86Sr have been
calculated at 200 Ma which is the 40Ar-39Ar age for these intrusive rocks (e.g. Marzoli
et al., 1999) (Fig. 5.5). The initial εNd values for the Palisade Sill basalts and gabbros
range from -0.64 to -3.8 whereas the sandstones and chilled margin basalts have a
much more negative εNd(I) range of -6.3 to -12.3. The initial 87Sr/86Sr values of basalts
from the Palisade Sill range from 0.70668 to 0.71037, with the sandstones and chilled
margins displaying the most radiogenic 87Sr/86Sr(I) from 0.70736 to 0.72267. In figure
5.5 all basalts, sandstones, and chilled margins of this study have been compared to
the Central Atlantic MORB (Janney and Castillo, 1996) as well as CAMP related
intrusives from Liberia (Dupuy et al., 1988), eastern North America (Peagram, 1990;
(Pegram, 1990; Heatherington and Mueller, 1999), Spain (Alibert, 1985; Cebria et al.,
2003), Guinea and Guyana (Deckart et al., 1997), Brazil (DeMin et al., 2003), France
(Jourdan et al., 2003), and Nova Scotia (Pe-Piper and Reynolds, 2000).
In the Nd-Sr isotopic correlation (Fig. 5.5) the Palisade Sill basalts show
affinity with the low-Ti CAMP tholeiites and lie in the enriched quadrant with
negative εNd and radiogenic 87Sr/86Sr in contrast to the high-Ti CAMP tholeiites
which have a smaller isotopic range and lie in the quadrant of mantle derived rocks. A
mixing is constructed with the Central Atlantic Plume (CAP) (Cebria et al., 2003) as
231
the uncontaminated plume end member and sandstone from the Palisade Sill region as
the contaminant end member.
Initial 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb of the Palisade Sill basalts and
gabbros at 200 Ma have ranges of 18.11-18.69, 15.57-15.69, and 37.74-39.82 and the
sandstones and chilled margin have ranges of 18.45-18.67, 15.62-15.77, and 38.06-
39.78 respectively as reported in Table 5.3. Pb-Pb isotopic ratios of these rock are
plotted in figure 5.6 along with Central Atlantic MORB (Janney and Castillo, 1996)
as well as CAMP related intrusives from eastern North America (Pegram, 1990;
Heatherington and Mueller, 1999), Spain (Alibert, 1985; Cebria et al., 2003), Guinea
and Guyana (Deckart et al., 1997), France (Jourdan et al., 2003), Nova Scotia (Pe-
Piper and Reynolds, 2000), and mafic and felsic gneisses from Honeybrook Upland,
Pennsylvania (Sinha et al., 1996). Various continental crustal and mantle reservoirs as
well as the Northern Hemisphere Reference Line (NHRL) are also plotted in the Pb-
Pb plots for reference (Zartman and Doe, 1981; Hart and Zindler, 1989). Notice the
correspondence of the Palisade Sill basalts and gabbros with the other low-Ti CAMP
related tholeiites in figure 5.6b.
Initial εNd and 87Sr/86Sr vs. 206Pb/204Pb(I) at 200 Ma for all the rocks of this
study are shown in figure 5.7a and b respectively and compared with relevant CAMP
related basalts from eastern North America (Pegram, 1990; Heatherington and
Mueller, 1999), Spain (Alibert, 1985; Cebria et al., 2003), Guinea and Guyana
(Deckart et al., 1997), France (Jourdan et al., 2003), Nova Scotia (Pe-Piper and
Reynolds, 2000), Central Atlantic MORB (Janney and Castillo, 1996), and various
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Table 5.2. Initial Nd-Sr-Pb-isotopic data of the Palisade Sill samples of this study
corrected to 200 Ma
147Sm/
144Nd 143Nd/
144Nd(0) 143Nd/
144Nd(I) εNd(I)
87Rb/ 86Sr
87Sr/86Sr(0) 87Sr/86Sr(I)
PS-01 Basalts 0.16 0.512492 0.512291 -1.8 0.51 0.708061 0.706614 PS-02 0.16 0.512453 0.512242 -2.7 0.34 0.706832 0.705857 PS-03 0.16 0.512509 0.512299 -1.6 0.38 0.707276 0.706186 PS-05 0.16 0.512547 0.512331 -1.0 0.47 0.707573 0.706226 PS-07 0.16 0.512388 0.512177 -4.0 0.33 0.707229 0.706299 PS-08 0.17 0.512533 0.512316 -1.3 0.44 0.707887 0.706639
PS-10A 0.15 0.512472 0.512276 -2.0 0.31 0.707238 0.706356
PS-10B 0.15 0.512522 0.512322 -1.2 0.43 0.706683 0.705454
PS-11 0.15 0.512500 0.512305 -1.5 0.44 0.708068 0.706825 PS-12 0.14 0.512488 0.512303 -1.5 0.29 0.706820 0.706002 PS-13 0.14 0.512506 0.512322 -1.12 0.83 0.710371 0.708008 PS-14 0.15 0.512489 0.512296 -1.7 0.39 0.707397 0.706301 PS-17 0.14 0.512467 0.512286 -1.7 0.36 0.706937 0.705926 PS-20 0.15 0.512509 0.512311 -1.4 0.29 0.706914 0.706096 PS-21 0.15 0.512474 0.512275 -2.1 0.18 0.706960 0.706454 PS-23 0.16 0.512451 0.512242 -2.7 0.36 0.707408 0.706381 PS-24 0.16 0.512562 0.512348 -0.6 0.21 0.706817 0.706219 PS-04 Chilled
Margins 0.14 0.512044 0.511861 -10.1 2.77 0.721529 0.713646
PS-15 0.11 0.512108 0.511968 -8.1 2.27 0.722667 0.716224 PS-22 0.60 0.512537 0.511754 -12.3 0.34 0.707355 0.706378 PS-06 Sand-
stones 0.13 0.512104 0.511933 -8.7 0.91 0.714436 0.711853
PS-09 0.12 0.512105 0.511953 -8.4 2.36 0.721286 0.714571 PS-16 0.13 0.512138 0.511966 -8.1 0.55 0.714879 0.713314 PS-19 0.21 0.512338 0.512061 -6.3 1.08 0.717792 0.714725
233
Table 5.2 continued.
206Pb/ 204Pb(0)
207Pb/ 204Pb(0)
208Pb/ 204Pb(0)
238U/ 204Pb
235U/ 204Pb
232Th/ 204Pb
206Pb/ 204Pb(I)
207Pb/ 204Pb(I)
208Pb/ 204Pb(I)
PS-01 18.56 15.62 38.50 9.23 0.07 27.99 18.27 15.61 37.96 PS-02 18.55 15.70 40.20 5.77 0.04 25.03 18.37 15.69 39.72 PS-03 18.33 15.57 38.14 2.66 0.02 11.63 18.25 15.57 37.92 PS-05 18.33 15.64 38.26 2.13 0.02 8.88 18.26 15.64 38.09 PS-07 18.50 15.63 38.31 2.29 0.02 7.67 18.42 15.62 38.16 PS-08 18.39 15.65 39.15 2.63 0.02 12.04 18.31 15.64 38.92 PS-10A 18.52 15.63 40.05 5.49 0.04 24.06 18.35 15.62 39.59 PS-10B 18.42 15.65 39.99 2.12 0.02 8.29 18.35 15.65 39.83 PS-11 18.48 15.67 38.52 3.61 0.03 15.68 18.37 15.66 38.22 PS-12 18.24 15.62 38.14 1.25 0.01 3.36 18.20 15.62 38.08 PS-13 18.25 15.63 38.13 0.70 0.01 2.90 18.23 15.62 38.08 PS-14 18.43 15.69 38.53 4.53 0.03 19.36 18.29 15.68 38.16 PS-17 18.25 15.63 38.17 1.49 0.01 4.89 18.21 15.62 38.07 PS-20 18.22 15.58 38.03 3.30 0.02 14.85 18.11 15.57 37.74 PS-21 18.22 15.59 38.04 0.95 0.01 4.29 18.19 15.58 37.96 PS-23 18.73 15.64 38.37 1.22 0.01 5.10 18.69 15.64 38.27 PS-24 18.44 15.65 38.45 4.84 0.04 21.84 18.28 15.64 38.03 PS-04 18.76 15.65 39.00 5.54 0.04 34.99 18.59 15.64 38.33 PS-15 18.86 15.64 38.60 5.96 0.04 23.38 18.67 15.63 38.15 PS-22 18.54 15.62 38.44 2.71 0.02 11.83 18.45 15.62 38.22 PS-06 18.84 15.72 20.26 7.98 0.06 43.28 18.58 15.71 39.43 PS-09 18.85 15.78 40.18 8.19 0.06 87.68 18.59 15.77 38.50 PS-16 18.59 15.66 39.90 2.42 0.02 6.50 18.51 15.65 39.78 PS-19 18.61 15.65 38.72 4.19 0.03 34.37 18.48 15.64 38.06
234
mantle reservoirs (Zartman and Doe, 1981; Hart and Zindler, 1989). Here again, the
Palisade Sill basalts show affinity towards other CAMP related low-Ti rocks.
5.4. Discussion of Palisade Sill Geochemistry compared to existing high and low-
Ti CAMP from literature
The participation of a mantle plume as the main cause of the CAMP is a
matter of intense debate and there are many arguments that can be put forward to
either support (e.g. Ernst et al., 1995; Wilson, 1997) or reject (e.g. Bertrand, 1991;
McHone, 2000) the participation of a Central Atlantic Plume (CAP) in the
petrogenesis of the primary magmas. In this section we discuss the geochemical
results presented in section 5.3 and attempt to identify either a plume or a
subcontinental mantle lithospheric source for the Palisade Sill basalts as well as other
low-Ti CAMP tholeiites.
5.4.1. Discussion of Trace Element Geochemistry
The Chondrite normalized REE patterns of the basalts and the gabbros from
the Palisade Sill display only minor variability along the entire Sill (Fig. 5.2a). There
is slight LREE enrichment in these rocks which is typical of both high and low-Ti
CAMP tholeiites (e.g. Alibert, 1985; Dupuy et al., 1988; Bertrand, 1991; DeMin et
al., 2003). In particular, the REE patterns of the Palisade Sill basalts strongly
resemble the low-Ti tholeiites from Europe (Jourdan et al., 2003) and Brazil (DeMin
et al., 2003). The Palisade Sill basalts as well as other high and low-Ti CAMP
235
tholeiites are distinctly different from Central Atlantic MORB (Janney and Castillo,
1996) in their REE patterns and are likely derived from an enriched mantle. In
contrast to the restricted REE patterns of the Palisade basalts, the sandstones and
chilled margins have a much wider range of REEs (Fig. 5.2b) with a strong negative
Eu anomaly that is characteristic of continental crust. The sandstones of this study are
the likely contaminants of the Palisade Sill basalts. The chilled margins that formed
by the sudden cooling of the intrusive lava with the surrounding rock have REE
patterns similar to the sandstones (Fig. 5.2b) and likely experience nearly 100%
contamination by the sandstones.
On the multi-element primitive mantle normalized diagram the Palisade
basalts are flat with the exception of low Ba and very high Pb (Fig. 5.3a). There are
very small negative Nb-Ta and positive Zr-Ha anomalies in a few samples. Tholeiites
from French Guyana (Nomade et al., 2002) and Spain (Cebria et al., 2003) also show
similar patterns. The nearly flat HREE pattern implies an absence of garnet in the
source; the depth of melting for these rocks may have been in the field of stability of
spinel peridotite. These rocks were possibly derived from a mantle ~20-50 times
more enriched than the primitive mantle. The sandstones and chilled margins have
low Ba, Sr and Eu, negative Nb-Ta anomalies, and high Pb when normalized to the
primitive mantle (Fig. 5.3b). Notice that both Palisade Sill basalts as well as
sandstones show the enrichments and depletions in the same element, but the
magnitude of enrichment and depletion is higher in the Sandstones compared to the
Palisade basalts. This indicates that the Palisade Sill lavas to be contaminated by
236
these sandstones prior to or during emplacement. The chilled margins experienced
very large degrees of contamination by the sandstones and hence mimic their trace
element patterns (Figs. 5.2, 5.3).
Most continental flood basalts (CFBs) are characterized by both high and low-
Ti basalts (Saunders et al., 1992; Hawkesworth et al., 1999). The Palisade Sill basalts
appear to be derived from a source region similar to other low-Ti CAMP (e.g.
Pegram, 1990; Deckart et al., 1997; Jourdan et al., 2003) as well as other continental
flood basalts such as the Parana and (Hawkesworth et al., 1999; Nomade et al., 2002)
and in figure 5.4a. These basalts have similar La/Ba ratios and slightly higher La/Nb
ratios compared to ocean island basalts. They are distinctly different from the high-Ti
CAMP tholeiites from Guyana (Deckart et al., 1997) which lie entirely within the
ocean island basalt field and have relatively higher La/Ba ratios.
The systematic differences between the high and low-Ti CAMP tholeiites may
indicate that these two types of tholeiites have different mantle sources (e.g. Cebria et
al., 2003). However, it is also possible that both high and low-Ti tholeiites are derived
from an enriched mantle and the low-Ti tholeiites have been contaminated by the
continental lithosphere whereas the high-Ti tholeiites lie close to the primitive mantle
and may or may not have asthenospheric contamination. Another possibility is that
the high and low-Ti lavas of the CAMP provinces may have been derived by different
depths and degrees of melting of the same mantle.
Studies of Nb-Ta variation in MORBs, komatiites, depleted mantle xenoliths,
and chondritic meteorites (Jochum et al., 1986; Jochum et al., 1989) have suggested
237
that the Nb/Ta of the Earth’s mantle has a chondritic value of 17.5 and that there is no
significant fractionation of these elements, at least at large degrees of partial melting.
However, experimental studies (Green and Pearson, 1987; Green et al., 1989; Green,
1995) of the partitioning of these elements between various melt compositions and
potential residual phases in the mantle suggest that in certain circumstances this pair
of geochemically similar elements may be fractionated. These effects have been
modeled assuming primitive Nb and Ta concentrations, and the calculated likely
maximum variation of Nb concentration with Nb/Ta in the melt and the residue at
various degrees of batch meting in figure 5.4b.
At very low degrees of melting (<1%) Nb/Ta values may be similar to
leucites, but calculated Nb concentrations are much higher, and Zr/Nb values are
much lower than for leucites (Stolz et al., 1996). However, if a slightly depleted
mantle source is used as the starting composition, the calculated Nb/Ta values for the
melt are much lower (~14) due to the significant reduction of Nb/Ta in the residue
after removal of a very small melt fractions. Hence from figure 5.4b we can infer that
the low-Ti Palisade Sill basalts were derived from ~15% melting of a slightly
depleted peridotite. The high-Ti tholeiites may have been products of the initial small
melt fractions that left behind a slightly depleted peridotite source for the low-Ti
CAMP tholeiites.
The Palisade Sill basalts and most of the CAMP investigated throughout the
four circum-Atlantic continents are typically tholeiitic low-Ti continental flood
basalts, which differ fundamentally from MORB by higher concentrations of LREE
238
(Figs. 5.3-5.4) (e.g. Bertrand et al., 1982; Alibert, 1985; Dupuy et al., 1988; Bertrand,
1991; Puffer, 2002; Cebria et al., 2003; DeMin et al., 2003; Jourdan et al., 2003). On
the basis of our trace element data discussed above, the low-Ti Palisade Sill basalts
appear to have been derived by melting of a spinel lherzolite and are likely
contaminated by the continental crust represented by the sandstones and chilled
margin basalts of this study.
5.4.2. Discussion of Nd-Sr-Pb Geochemistry
Despite their apparent similarity in the normalized trace element diagrams, the
high and low-Ti CAMP magmatic rocks exhibit two distinct ranges of Sr-Nd
compositions. The enriched Nd-Sr composition of the Palisade basalts suggests
contamination by a more radiogenic component. Several studies have explained the
Nd-Sr isotopic composition of the low-Ti CAMP rocks by fractional crystallization of
an enriched lithospheric mantle coupled with assimilation of lower crustal granulites
(e.g. Cebria et al., 2003). One of the key questions concerning the Palisade Sill and in
general the CAMP giant dike swarms is to determine if the dikes are a result of lateral
injection of magmas radiating from a CAP or if they originate from other mantle
sources. Dike emplacement models confirm that magmas can travel laterally long
distance (>3000km) from their focal point (Elliot et al., 1999). Although the larger
number of CAMP related magmas with radiogenic Sr and negative εNd may suggest
the mantle lithosphere to be the main source of these rocks, the intermediate
composition of these low-Ti tholeiites between CAP and crust-like compositions
239
points towards continentally contaminated plume derived melts for these dikes and
sills.
The presence of a plume related mantle source component in the CAMP is
supported by data published for the oldest (160-120Ma) Atlantic Oceanic crust
(Janney and Castillo, 1996), which suggest the involvement of a plume type mantle
component during the early stages of the Central Atlantic opening. Similarly, the
primitive olivine dolerites and lamprophyre dikes from Nova Scotia, Canada, also
suggest the presence of plume sources in CAMP (Pe-Piper and Reynolds, 2000).
Furthermore, as we have shown in our trace element data the Palisade Sill basalts as
well as other low-Ti CAMP data support the participation of an OIB type mantle
source (Figs. 5.2, 5.3).
We use the age-corrected (200Ma) isotopic composition of the Central
Atlantic Plume (CAP) (Cebria et al., 2003) as the starting composition of the Palisade
Sill basalts. As observed from Fig. 5.5 the assumed CAP derived tholeiites
assimilated ~ 10-20% continental crust, as represented by the sandstone, to produce
the Palisade Sill basalts. Also note the absence of a Central Atlantic MORB like
contaminant in both the high and low-Ti tholeiites of the CAMP as well as the
Palisade Sill basalts of this study. Since the CAMP consists of intrusive rocks,
assimilation of mid-continental crust during emplacement and ponding of the magmas
in the crust is expected.
207Pb/204Pb(I) and 208Pb/204Pb(I) for the CAMP magmatic rocks correlate
positively with 206Pb/204Pb(I) defining a tight linear array subparallel to the NHRL but
240
displaced towards higher 207Pb/204Pb and 208Pb/204Pb (Fig. 5.6). Taken as a whole, all
the low-Ti CAMP data collectively define a single positive correlation displaced
vertically above the field of MORB and resembles OIB from localities such as
Gough, Reunion, and French Polynesia (Pegram, 1990, and references therein). Such
207Pb/204Pb traits are characteristics of the majority of continental flood basalts as
well.
The Palisade Sill basalts overlap almost entirely with other low-Ti tholeiites
associated with the CAMP (Fig. 5.6). Collectively the Palisade Sill basalts and low-Ti
CAMP tholeiites trend towards EM-II which is considered to be middle or upper
continental crust (Fig. 5.6a). There is no correspondence of these low-Ti lavas with
either the average lower crust or with EM-I which represents sublithopheric mantle
compositions. There is also no correspondence of these data with Grenville age
granulites from Honeybrooks Creek, Pennsylvania (Sinha et al., 1996), which is
representative of the lower crust-mantle lithosphere. Hence it is likely that the low-Ti
CAMP basalts were derived from a mantle plume and assimilated continental crust
during emplacement. The overlap of the high and low-Ti CAMP related lavas with
the Central Atlantic MORB in their Pb-Pb correlation (Fig. 5.6) is not unexpected as
the oldest Atlantic oceanic crust has experienced widespread plume contamination
(Janney and Castillo, 1996). The absence of a depleted asthenosphere for the CAMP
rocks and specifically for the Palisade Sill basalts is confirmed by their normalized
trace element patterns (Fig. 5.2, 5.3) as discussed previously.
241
Figure 5.5. Initial εNd vs. 87Sr/86Sr at 200 Ma for the Palisade sill basalts and gabbros
compared to global CAMP data and the Central Atlantic MORB (Janney and Castilo,
2001). All CAMP data from literature are corrected to 200Ma. References for CAMP
data from literature are as follows: Liberia (Dupey et al., 1988); eastern North
America (Peagram, 1990; Heatherington and Mueller, 1991); Spain (Alibert, 1985;
Cebria et al., 2003); Guinea and Guyana (Deckart et al., 2005); Brazil (DeMin et al.,
2003); France (Jourdan et al., 2003); Nova Scotia (Pe-Piper and Reynolds, 2000).
242
Figure 5.6. Initial 208Pb/204Pb(I) vs. 206Pb/204Pb(I) and (b) 207Pb/204Pb(I) vs. 206Pb/204Pb(I)
plots of the Palisade sill basalts and gabbros compared to global CAMP data and
Central Atlantic MORB. Grenville granulites are from Sinha et al., (1996). Also
shown are the domains of the BSE – Bulk silicate Earth; NHRL – Northern
Hemispheric Reference Line; and DM – Depleted mantle (Hart and Zindler, 1989);
Upper crust and lower continental crust (Zartman and Doe, 1981). All other data
sources and symbols are as in figure 5.5.
243
The relatively small variation in Nd-Sr-Pb isotopic compositions of the
Palisade basalts suggests that the amount of crustal material assimilated was not very
significant (Fig. 6.7). The trend of the Palisade basalts towards the sandstones, chilled
margin basalts and EM-II is clear in the variation of Nd-Pb (Fig. 6.7a). The absence
of an EM-I or a Central Atlantic MORB like component in the Palisade lavas as well
as the low-Ti CAMP basalts in Nb-Pb-Sr variations (Fig. 6.7) re-emphasize our
inference that these rocks are likely derived from a plume source that was
contaminated by the continental crust, rather than my the melting of a heterogeneous
mantle lithosphere due to plume impingement. The overall similarity of CAMP
tholeiites has been used to support the involvement of a common mantle source
(Marzoli et al., 1999). The inferences made for these dikes and Sills, when considered
individually, imply the contamination of the original magma by the local crustal
component (e.g. Bertrand et al., 1982; Dunn et al., 1998).
Geochemical characteristics common to all low-Ti CAMP related tholeiites
including the Palisade Sill basalts are slight LREE enrichment (Fig. 5.2, 5.3),
enriched Nd-Sr compositions, and positive correlation in the Pb-Pb plot (Fig. 5.6)
subparallel to the NHRL. These trace element and Nd-Sr-Pb isotopic signatures
cannot unequivocally discriminate between a heterogeneous subcontinental
lithosphere and a plume source contaminated by the continental crust. However, the
absence of any EM-I component in any of the low-Ti CAMP along with presence of
an EM-II like contaminant in indicative of the latter. Also, major element chemistry
of the Palisade Sill rocks indicate 50-55 weight % SiO2 in these rocks
244
Figure 5.7. (a) Initial εNd vs. 206Pb/204Pb(I) and (b) 87Sr/86Sr(I) vs. 206Pb/204Pb(I) for the
Palisade sill basalts and gabbros compared to global CAMP data and Central Atlantic
MORB. Data sources and symbols are as in figures 5.5 and 5.6.
245
(e.g. Shirley, 1987; Husch, 1992; Puffer, 2002) indicating contamination and
assimilation of a silica rich continental crust rather than derivation from a silica poor
mantle lithosphere.
5.5. Geodynamic implications of our geochemical data for a plume source for the
Palisade Sill basalts and other high and low-Ti CAMP lavas
The relative enrichment of these elements in CAMP has been long debated
and explained by different mechanisms: (1) derivation from an enriched
subcontinental lithospheric mantle source, with no or limited crustal contamination
during magma ascent (e.g. Bertrand et al., 1982; Alibert, 1985; Dupuy et al., 1988;
Pegram, 1990; Bertrand, 1991; Heatherington and Mueller, 1999; DeMin et al.,
2003); (2) derivation from an asthenospheric MORB-like source with a more
significant crustal contamination (Dupuy and Dostal, 1984; Dostal and Durning,
1998); and (3) derivation from an incipient plume head (White and McKenzie, 1989;
Hill, 1991; Oyarzun et al., 1997; Wilson, 1997; Courtillot et al., 1999; Ernst and
Buchan, 2002; Morgan, 19893). With respect to (3) several authors discuss the
geochemical composition of the possibly involved magmas (Oliviera et al., 1990;
Janney and Castillo, 1996; Cebria et al., 2003).
Our geochemical data show no evidence of a depleted asthenospheric MORB
source for the CAMP derived dikes and sills. Derivation from an enriched
subcontinental lithosphere requires lithospheric thinning to drive decompressional
melting (White and McKenzie, 1989), which is unlikely to produce effusive, rapid
246
volcanism. In plume models, mantle material originates at or near the core-mantle
boundary and rises buoyantly through the mantle (e.g. Hill, 1991), spreading laterally
as the plume head encounters the thin eroded lithosphere and producing voluminous
melt. Is it possible that all the CAMP (6000 km) is a result of a single process:
melting of a heterogeneous lithospheric mantle initiated by plume impingement? A
more simple answer is to attribute the CAMP to a hotspot system. Geochronological
data show that most parts of the investigated CAMP were active in the early Jurassic
(~200Ma) (Deckart et al., 1997; Marzoli et al., 1999; Hames et al., 2000). This time
coincidence of the CAMP age argues for one plume for all the CAMP magmatism.
The brief and extremely widespread tholeiitic magmatism associated with the CAMP
implies that an anomalously hot mantle extended over a very wide area and melted
extensively. In general, our data and previous geochemical and geochronologial data
on CAMP are consistent with models that suggest that an upwelling plume head
separated from the plume tail (Leitch et al., 1998) and that the plume material spread
over a very large area by ambient mantle flux (Wilson, 1997).
The CAMP can be compared to the East African Rift System (EARS) which is
a classic example of ongoing continental rifting and provides an excellent framework
to investigate magmatism in an extensional setting. The EARS rift system extends
over 4000 km from the Red Sea in the north to Mozambique in the south representing
a ∼150 km wide zone of NW–SE trending extension (Chakrabarti et al., 2009). The
seismically and volcanically active EARS is the youngest mantle plume province
worldwide, with one or more upwellings impinging on thick cratonic lithosphere
247
since ~45 Ma (e.g. WoldeGabriel et al., 1990; Furman et al., 2006) caused by an
anomalously hot asthenosphere (Ebinger and Furman, 2003; Furman et al., 2004).
Discrete rifting episodes have recently been observed in the Afar triple junction in
Ethiopia. In this region batches of molten mantle rocks have risen into cracks and
fractures to form long, thin vertical sheets of new crust in the form of dikes, often
feeding surface eruptions of basalts. These dikes serve to transport melt percolating
upward from mantle source zones and also had accumulated in magma chambers or
thin horizontal magma sheets within the crust. The dikes along with faults constitute
plate boundary separation within the crust in this region (Barberi and Varet, 1977;
Hayward and Ebinger, 1996). Tomography and seismic images (Nyblade and
Robinson, 1994; Weeraratne et al., 2003), numerical models (Beutel et al., 2010), and
geochemical data (Marty et al., 1993; Furman et al., 2006; Pik et al., 2006;
Chakrabarti et al., 2009) collectively support a single plume with multiple stems
originating in this African super-plume.
The CAMP was derived from a compositionally heterogeneous super-plume
similar to the present day EARS, with multiple stems acting as feeders over a large
aerial extent.
5.6. Conclusions
In this study we have correlated the Palisade Sill basalts and gabbros to other
low-Ti CAMP related magmatism from eastern North America (Pegram, 1990;
Heatherington and Mueller, 1999), South America (DeMin et al., 2003), Europe
248
(Alibert, 1985; Cebria et al., 2003; Jourdan et al., 2003), West Africa (Deckart et al.,
1997), and Canada (Pe-Piper and Reynolds, 2000) by their geochemical and Nd-Sr-
Pb isotopic signatures. When compared to the small volume high-Ti CAMP related
magmatism, the low-Ti CAMP lavas have a more radiogenic Sr and less radiogenic
Nd component as well as higher values of 207PPb and 208Pb. Based on the
geochemical data presented in this study the Palisade Sill basalts were derived from a
slightly enriched OIB-like mantle source (Fig. 5.2, 5.3, 5.4a). Further, these rocks
were derived from ~15% (Fig. 5.4b) melting of a slightly depleted spinel peridotite
(Fig. 5.4c). Since other low-Ti CAMP lavas have similar geochemistry as well as
eruption ages as the Palisade Sill basalts of this study, it is safe to assume the same
source for these tholeiites across the four continents where they are emplaced.
Although the collective trace element and Nd-Sr-Pb isotopic signatures of
low-Ti CAMP magmatism cannot unequivocally discriminate between a
heterogeneous subcontinental lithosphere and a plume source contaminated by the
continental lithosphere, we argue for the latter based on the short time of eruption of
all these lavas as well as the absence of an EM-I like component in the low-Ti lavas.
We suggest that the CAMP was derived from a compositionally
heterogeneous super-plume similar to the present day East African Rift System, with
multiple stems acting as feeders over a large aerial extent. Compositional
heterogeneity may have been caused due to variability in lithosphere-asthenosphere
boundary, or due to different degrees and depths of melting as has been suggested for
the heterogeneous basalts derived from the EARS (Chakrabarti et al., 2009).
249
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