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The Graduate School
2006
Geochemical and Geochronological
Investigations in the Southern Appalachians,
Southern Rocky Mountains and Deccan
Traps.
Reshmi Das
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THE FLORIDA STATE UNIVERSITY
COLLEGE OF ARTS AND SCIENCES
GEOCHEMICAL AND GEOCHRONOLOGICAL INVESTIGATIONS IN THE
SOUTHERN APPALACHIANS, SOUTHERN ROCKY MOUNTAINS AND
DECCAN TRAPS.
By
RESHMI DAS
A Dissertation submitted to the
Department of Geological Sciences
in partial fulfillment of the
requirements for the degree of
Doctor of Philosophy
Degree Awarded:
Fall Semester, 2006
The members of the Committee approve the Dissertation of Reshmi Das defended on 1st
November, 2006.
____________________________________
A. Leroy Odom
Professor Directing Dissertation
____________________________________
Jeffrey Chanton
Outside Committee Member
____________________________________
Stephen A. Kish
Committee Member
____________________________________
Vincent J. M. Salters
Committee Member
____________________________________
James F. Tull
Committee Member
Approved:
_____________________________________________
Professor A. Leroy Odom, Chair, Geological Sciences
The Office of Graduate Studies has verified and approved the above named committee
members.
ii
To
my first geology teacher,
Dr. Mohan Chand Baral,
who
made me fall in love with Geology.
iii
ACKNOWLEDGEMENTS
Many people have contributed to the making of this dissertation. First and foremost I
would like to acknowledge Prof. Leroy Odom, my advisor who not only supported me
academically and financially for five years but also lent personal support whenever
required. This dissertation would have been impossible without him.
Prof. James Tull supervised my project on the southern Appalachian and words are
inadequate to express his contribution in development of this dissertation.
Prof. Stephen Kish always had an answer for my most difficult question and Prof.
Vincent Salters helped me with the modeling part of the chemical data.
Prof. Munir Humayun, Prof. Tapas Bhattacharyya and Dr. Michael Bizimis, though not a
part of my committee, supervised portions of my research. I am indebted to them for
their time, technical support and patience.
The field work required for this dissertation was partly funded by EDMAP component of
the National Geologic Mapping Act 2002.
The Isotope Geochemistry Division at the National High Magnetic Field Laboratory has
been a wonderful place to work, and I would like to thank Ted Zateslo and Afi SachiKocher for their technical support.
I would also like to thank all of the geochemistry graduate students for the long chats, the
little favors, and camaraderie along the way.
My parents made the greatest sacrifice of their life by letting their only child go ahead
with her life and providing the best educational opportunity.
Last but not the least is my best friend and husband Subhajit who inspired me to continue
my career in academia. His extremely active academic and intellectual presence kept me
going and will continue to do so.
iv
TABLE OF CONTENTS
List of Tables
................................................................................................
List of Figures
................................................................................................
Abstract
......................................................................................................
vii
viii
x
1. Geochemical and Geochronological Constraints on the Origin and Evolution of the Eastern
Blue Ridge, Southern Appalachians. ..........................................
1
1.1 Introduction...........................................................................................
1
1.2 Regional Overview ...............................................................................
3
1.3 Study Area and General Description of Lithotectonic Units................
4
1.4 Purpose of the Study .............................................................................
7
1.5 Analytical Techniques ..........................................................................
8
1.6 Results ................................................................................................
10
1.7 Nd Model Age and Detrital Zircon Ages of the Metasediments of the Eastern Blue
Ridge ..........................................................................................
14
1.8 Mulberry Rock Gneiss ..........................................................................
15
1.9 Discussion .............................................................................................
16
2. Kilometers Scale Strontium Isotopic Homogenization During Metamorphism: A Case Study
in the Tres Piedras Granite, New Mexico....................................
62
2.1 Introduction...........................................................................................
2.2 General Geology ...................................................................................
2.3 Tres Piedras Granite..............................................................................
2.4 Previous Work ......................................................................................
2.5 Results ................................................................................................
2.6 Discussion .............................................................................................
2.7 Conclusion ............................................................................................
62
63
64
65
69
70
74
3. Trace Element and Lead Isotopic Studies of the Kutch Volcanics of Northwest India
................................................................................................
87
3.1 Introduction...........................................................................................
3.2 Geological Setting.................................................................................
3.3 Deccan Stratigraphy..............................................................................
87
88
89
3.4 Previous Work-Sr and Nd Isotopic Data ..............................................
3.5 Analytical Technique ............................................................................
3.6 Results ................................................................................................
3.7 Comparison with the DVP ....................................................................
3.8 Discussion .............................................................................................
3.9 Conclusion ............................................................................................
APPENDICES
90
91
91
93
94
95
................................................................................................
107
A Analytical Techniques ..........................................................................
B Sample Locations..................................................................................
C Photomicrographs .................................................................................
107
118
120
REFERENCES
................................................................................................
122
BIOGRAPHICAL SKETCH ..............................................................................
149
vi
LIST OF TABLES
Table 1.1: Major Oxide concentration in weight percent. ...................................
26
Table 1.2: Trace Element and REE concentration (in ppm) of the Pumpkinvine Amphibolites
(PCF) and the Galts Ferry Gneiss (GFG) ............................................................
27
Table 1.3: Trace Element and REE concentration (in ppm) of the Hillabee Greenstones (HG) and
Hillabee Dacite (dacite) .......................................................................................
28
Table 1.4: Sr and Nd Isotopic Data......................................................................
29
Table 1.5: Rb-Sr isotopic data for whole rock samples.......................................
30
Table 1.6: Nd Model Age of metasediments from eastern Blue Ridge ...............
30
Table 1.7: U-Pb analysis of Galts Ferry Gneiss zircons by LA-MS-ICPMS ......
31
Table 1.8: U-Pb analysis of detrital zircons from meta-sandstone by LA-MS-ICPMS 33
Table 1.9: U-Pb analysis of Mulberry Rock Gneiss zircons by LA-MS-ICPMS
35
Table 1.10: REE concentrations of eastern Blue Ridge metasediments..............
36
Table 2.1: Regional Geologic history of the north-central New Mexico ............
67
Table 2.2: Stratigraphic nomenclature and lithologic description of supracrustal Proterozoic
rocks of New Mexico...........................................................................................
68
Table 2.3: Chemical analysis, norm and modes of the Tres Piedras Granite ......
75
Table 2.4: Rb-Sr whole rock analysis of Tres Piedras Granite............................
76
Table 2.5: Rb-Sr analysis of the mineral phases of the Tres Piedras Granite......
77
Table 2.6: Tres Piedras Granite zircon analysis by LA-MC-ICPMS ..................
78
Table 3.1: Stratigraphic nomenclature and thickness of the southwestern Deccan Formations
................................................................................................
97
Table 3.2: Trace element and Pb-isotope results .................................................
98
Table 3.3: Elemental ratios of Re'union type source at 65 Ma ............................
94
vii
LIST OF FIGURES
Figure 1.1: Generalized geological map of southern Appalachians ...................
37
Figure 1.1A: Generalized geologic map of eastern Alabama and western Georgia Blue Ridge
terranes
................................................................................................
38
Figure 1.2: Bimodality of Hillabee Greenstone sequence and Pumpkinvine Creek Formation
................................................................................................
39
Figure 1.3: Total alkali vs. silica diagram for the Hillabee Greenstone sequence and
Pumpkinvine Creek Formation ...........................................................................
40
Figure 1.4A: Hillabee dacite and GFG classification using Shand's index ........
41
Figure 1.4B: HG and PCF amphibolite plotted on AFM diagram ......................
41
Figure 1.5: Ti vs Zr plot of greenstones- dacites and PCF amphibolite - GFG ..
42
Figure 1.6: Co-variation diagrams of relatively immobile elements for HG and PCF amphibolites
................................................................................................
43
Figure 1.7: Tectonic discrimination diagrams for the PCF amphibolites and HG
44
Figure 1.8: La/10-Y/15-Nb/8 diagram for the PCF amphibolite and HG ..........
45
Figure 1.9: Tectonic discrimination diagram for GFG and Hillabee dacite .......
46
Figure 1.10: Spider diagram plot of PCF amphibolite and HG ..........................
47
Figure 1:11 CI chondrite normalized plot of PCF amphibolite and HG ............
48
Figure 1.12: CI chondrite normalized plot of GFG and Hillabee dacite ............
49
Figure 1.13: Initial Sr vs. epsilon Nd isotopic plot .............................................
50
Figure 1.14: Rb-Sr isochron of GFG samples ....................................................
51
Figure 1.15: Two chemical groups of GFG defined by REE concentrations. ....
52
Figure 1.16: U-Pb ages of GFG zircons .............................................................
53
Figure 1.16A: CL images of GFG zircons ..........................................................
54
Figure 1.17: Nd isotopic evolution diagram of the eBR metasediments ............
55
viii
Figure 1.18: NASC normalized REE plot of the eBR metasediments ...............
56
Figure 1.19: U-Pb ages of the detrital zircons of eBR ........................................
57
Figure 1.20: Rb-Sr whole rock age of Mulberry Rock Gneiss ...........................
58
Figure 1.21: U-Pb ages of Mulberry Rock Gneiss zircons .................................
59
Figure 1.22: Epsilon Nd of Ordovician arc felsic magmas of the Appalachians.
60
Figure 1.23: Model for passive mechanism of lithosphere extension ................
61
Figure 1.24: Model for accretionary orogen in southern Appalachians during Ordovician
................................................................................................
61
Figure 2.1: Map of Proterozoic basement uplifts and associated major faults in north-central New
Mexico
................................................................................................
80
Figure 2.1a: Map of Tres Piedras Granite sampling locations ...........................
80
Figure 2.2: Whole rock Rb-Sr isochron of Tres Piedras Granite ........................
81
Figure 2.3: Rb-Sr mineral isochron of Tres Piedras Granite ..............................
83
Figure 2.4: U-Pb ages of Tres Piedras Granite zircons by LA-MC-ICPMS ......
85
Figure 2.5: LA-ICPMS measurement of concentration ratios across a feldspar grain from Tres
Piedras Granite ................................................................................................
86
Figure 3.1: Structural-tectonic features of southern Asia and Indian Ocean basin
99
Figure 3.2: Initial Nd vs. initial Sr isotope plot of the Kutch samples ...............
100
Figure 3.3: Trace element-REE plot of the Kutch basalts .................................
101
Figure 3.4: La/Nb vs Ce/Pb and La/Sm vs Sm/Yb in the Kutch basalts ............
102
Figure 3.5: Pb and Sr isotopic variation in Kutch Basalts ..................................
103
Figure 3.6: Melting model of garnet peridotite to produce the alkali basalts .....
104
Figure 3.7: Mixing model for the Kutch tholeiites to generate the trace element-REE pattern
................................................................................................
105
Figure 3.8: Mixing model for the Kutch tholeiites to generate the Sr and Nd isotope ratioserate
the trace element-REE pattern ............................................................................
106
ix
ABSTRACT
In the southernmost Appalachians, bimodal volcanics of Pumpkinvine Creek Formation (PCF)
and its proposed equivalent, the Hillabee Greenstone (HG) have indistinguishable ages (~460
Ma) and trace element-REE pattern similar to an arc/back-arc type setting. 143Nd values of
felsic members of the HG and PCF indicate involvement of Grenville crust during petrogenesis.
U-Pb dates (900-1500 Ma) of detrital zircons in PCF meta-sandstone cluster around 1100 Ma.
Nd-model ages of the Ashland-Wedowee Supergroup metasediments range between 943-1439
Ma and cluster around 1000 Ma. Rb-Sr whole rock and U-Pb zircon dates of the Mulberry Rock
Gneiss also demonstrate an Ordovician age (~460 Ma). It is concluded that the PCF-HG arc
formed on the Laurentian continental margin on Ashland-Wedowee sediments during
Ordovician and remained outboard of the continent until final closure during Alleghenian
orogeny.
Geochronological investigations of the Tres Piedras Granite of northcentral New Mexico have
revealed a sharp discordancy between Rb-Sr whole-rock and U-Pb zircon ages. Analyses of fifty
individual zircons (most concordant) by LA-MS-ICPMS yield a ~1730 Ma magmatic
crystallization age. Rb-Sr whole-rock isochron ages from separate localities are 1490+/-20 Ma
and 1497+/-42 Ma. Sphene/whole-rock/biotite isochron ages and initial 87Sr/86Sr ratios from
separate localities are indistinguishable from those of whole-rock isochrons. In both cases
feldspar plots above the isochrons and appeared to be an open system as evidenced by 4%
difference in 87Sr/86Sr in the core and rim of feldspar. Taken altogether, these data are interpreted
to reflect a large (kilometer) scale redistribution and rehomogenization of strontium isotopes
during an independently, well-documented metamorphic event in the region.
The geochemical character of the Kutch volcanics, northwest of Deccan Traps, India, have been
investigated in order the magma's origin. Sr-Nd-Pb isotopic ratios and trace element patterns
identifies three end members: Reunion plume-type alkali basalts, Mahabaleshwar-type alkali
basalts and crustally contaminated tholeiites. The first type of alkali basalts that can be generated
by very low degree of partial melting (1.6-1.8%) of Reunion plume like source at garnet stability
field; the tholeiites can be explained by crustal contamination of Indian-MORB like magma.
High 207Pb/204Pb (15.61-15.83) ratio of the tholeiites agrees well with the Pb-isotopes of local
Archean crust.
x
CHAPTER ONE
GEOCHEMICAL AND GEOCHRONOLOGICAL CONSTRAINTS ON THE ORIGIN AND
EVOLUTION OF THE EASTERN BLUE RIDGE, SOUTHERN APPALACHIANS.
1.1 Introduction
Forty years ago Tuzo Wilson (1966) proposed the theory of the Wilson Cycle by studying the
eastern margin of North America. At its simplest, the Wilson Cycle describes the opening and
closing of ocean basins - plates rift into pieces and diverge away from each other and new ocean
basins form in between, then there is a reversal in motion, convergence of the two rifted plates
followed by a plate collision, and mountain building. The rock records of the eastern margin of
North America preserves evidence of at least two complete Wilson cycles– starting with breakup
of the supercontinent Rodinia and ending with the assembly of the supercontinent Pangaea and
finally rifting of Pangea and opening of the modern Atlantic Ocean (Thomas, 2005).
The breakup and rifting of Rodinia is well documented in the stratigraphic rock records of the
eastern margin of North America that include synrift sedimentary and igneous rocks, post-rift
unconformity and early post-rift sedimentary strata. Isolation of rifted Laurentia was complete by
early Cambrian (ca. 530 Ma) (Thomas, 2005). The early extensions span between 760–650 Ma
(e.g., Aleinikoff et al., 1995; Hogan and Gilbert, 1998; Thomas et al., 2000; Cawood and
Nemchin, 2001; Owens and Tucker, 2003) followed by pervasive rifting (620–545 Ma) along
the eastern margin of Laurentia and terminating with the evidence of few late stage rifting of
microcontinents between 540–530 Ma.
The Appalachian mountain chain was created by broadly three Paleozoic orogenies after the
rifting of supercontinent Rodinia: the Ordovician-Silurian Taconic Orogeny, the Devonian
Acadian orogeny and the Pennsylvanian-Permian Alleghanian orogeny.
Evidence of deformation due to Taconic orogeny was first recognized in the Hudson valley of
New York and the affected areas stretch from western Newfoundland south to about New York
State. Deformation caused by Taconic Orogeny are also identified much farther south in Georgia
and Alabama. Taconic orogeny was dominated by A-type subduction against the Laurentian
margin resulting in the closure of the Proto Atlantic (also called the Iapetus Ocean). Island arcs
forming due to the subduction were obducted onto the Laurentian margin. Few examples include
Cowrock, Cartoogechaye and Tugaloo terranes in the southern Appalachians and the
Chopawamsic, Potomac and Bultimore terranes in the central and northern Appalachians
(Hatcher et al., 2006). The arc building and amalgamation produced numerous plutons
(Steltenpohl et al., 2005; Bream, 2003), penetrative deformation, and as high as granulite facies
metamorphism (Hatcher and Butler, 1979; Eckert et al., 1989; Moecher et al., 2004). Taconic
orogeny generated a huge amount of new crust that was amalgamated along the southeastern
margin of Laurentia.
The Acadian Orogeny occurred during the early Devonian Period. It was a continuing collision
to the island arc of the Taconic Orogeny (Hatcher, 1989). The Acadian Orogeny was primarily a
continent-continent collision between Laurentia (North America) and Baltica (Europe). The
1
Northern Iapetus Ocean was completely closed and the Acadian orogeny re-deformed some of
the rocks previously affected by the Taconic Orogeny, and produced a mountain belt in , Nova
Scotia, Newfoundland, Maine, New York, England, Ireland, Scandinavia and Greenland.that are
as high as the Himalayas The Acadian Orogeny of North America is equivalent to the
Caledonian Orogeny in Europe. Following the collision of Europe and North America, the newly
formed Appalachian Mountains started to shed tremendous amounts of sediment westward into
the foreland. The sediments were primarily conglomerates, quartz arenites, breccias, arkose
sandstones etc (Harrison, 2002, Hatcher, 2005). Near the mountains, the sedimentary rocks were
primarily terrestrial like river conglomerates and red shale, sands, etc. Westward, these rocks
undergo facies changes into black shales and beach sands (Baranoski and Riley, 1987). In
Pennsylvania and New York, the Catskill Delta represents the terrestrial-marine component (Van
Tassell, 1987). In other parts of North America, sediment erosion from the Appalachian
Mountains wasn’t a dominating event and the shallow marine seaways existed supporting large
amounts of carbonate sedimentation and oolites and reefs formation (Drivet and Mountjoy,
1997).
The Alleghenian Orogeny of the southern Appalachians was the third and last of the great
orogenies that affected eastern North America during the Paleozoic. The final collision of North
America with Africa, transported a huge composite crystalline thrust sheet – the Blue Ridge
Piedmont megathrust sheet that included pre-Alleghanian metamorphic rocks to at least 350 Km
inside the North American craton (Hatcher et al., 2006). The Alleghenian orogeny was
completed by 265 Ma with the formation of Pangea and completion of the first (Paleozoic)
Wilson cycle. The Alleghenian suture is marked by the boundary between the easternmost
subsurface Suwannee terrane and Carolina terrane. From the fossil assemblage of the cover
sequence (Gondwana basement covered by fossiliferous Ordovician to Devonian sandstones and
shales) it seems likely that the Suwanee originated as part of African Gondwanaland (Mueller et
al., 1994; Pojeta et al., 1976) and remained on the Laurentian side during breakup of Pangea.
The three episodes of orogeny are identified in the Appalachians but they are not substantiated
along the entire length of the mountain chain. (Moecher et al., 2004: Hatcher, 1978; Hatcher et
al., 1989). The southern Appalachian mountain chain records protracted continental orogeny
related to multiple extensional and contractional events. These events have been attributed to a
range of convergent tectonic mechanisms, including subduction, arc accretion, and continental
collision. Both direct and oblique convergence mechanisms have been proposed, and subduction
is interpreted to have been directed both toward and away from Laurentia, the core of North
America that existed at the time (Hatcher et al., 1987). The eastern Blue Ridge province was
affected by all of the major events that shaped the southern Appalachian orogen and is a key link
in elucidating the assembly of the Appalachian portion of the North American continent.
However, important aspects of the evolution of the eastern Blue Ridge remain obscure.
This work presents new data on zircon contained within the metasediments in the eastern Blue
Ridge of Alabama and Georgia. The geochemical and geochronological data have bearing upon
the formation, evolution and accretion of the most inboard ‘suspect’ terranes in the southern
Appalachians and will compare this sequence chemically with the Hillabee Greenstone of the
Talladega belt-western Blue Ridge. These data will have implication for the genesis of the
2
eastern Blue Ridge in the southern Appalachians and provide the implicit constraints on the
tectonic history of this region.
1.2 Regional Overview
The Blue Ridge province of the southern and central Appalachians includes low to high grade
metamorphic rocks that have been thrust northwestward over the unmetamorphosed sedimentary
rocks of the Valley and Ridge province. It is divided into eastern and western portions by the
east-dipping Hayesville and related faults in the northeast and by the Allatoona fault and Hollis
Line Fault in the southeast (Figs. 1.1, 1.1A). The post-metamorphic Allatoona fault and Hollis
Line fault places the structurally complex eastern Blue Ridge assemblage of chemically
immature clastic metasedimentary rocks, mafic to ultramafic bodies, variably deformed felsic
intrusive rocks and Grenville basement (eg.Tullahh Falls Formation) over the more mature
metasedimentary rocks and basement of the western Blue Ridge (Rankin, 1975; Hatcher, 1978).
Regionally the eastern Blue Ridge is a composite terrane that may include parts of Laurentian
outer margin cover sequence as well as accreted components of accretionary prism, ophiolitic
and island arc affinity. The eastern Blue Ridge (EBR) is generally similar lithologically to the
adjacent Piedmont. The Piedmont is also an exotic composite terrane that extends some 700 km
southwestward from the frame of the Sauratown Mountains window in North Carolina to the
Coastal Plain overlap in Alabama (Hatcher, 2002). The Brevard fault zone, that separates the
eastern Blue Ridge from the Piedmont, is a major structure with a protracted, polyphase history
of displacement (Hatcher, 1978). The western Blue Ridge is demonstrably part of Laurentia, but
the origin of the EBR is controversial. The presence of alpine peridotite, mafic-ultramafic
complexes has led some to suggest that the EBR may be an exotic, far-traveled terrane that was
accreted during Paleozoic orogeny, with the mafic rocks being remnants of the closed ocean
basin (for example, Horton, Drake, and Rankin, 1989; Willard and Adams, 1994; Zen, 1981;
Williams and Hatcher, 1983; Shaw and Wasserburg, 1984).
The Pine Mountain Window (Fig. 1.1) is exposed in the Piedmont on the Alabama promontory.
It exposes 1.1 Ga Laurentian Grenvillian basement of granulite- and upper-amphibolite-facies
granitic gneisses and a kyanite-sillimanite-grade platformal cover sequence (Hatcher, 2006).
The easternmost accreted terrane in the central and southern Appalachians is the Carolina
superterrane situated to the east of Piedmont (Fig. 1.1). During Rodinia rifting numerous
microcontinents were formed in the prevailing ocean between Laurentia and Gondwana,
proximal to Gondwana, loosely termed as Peri-Gondwanan terranes. The Carolina superterrane
was formed by amalgamation of a series of these microcontinents of peri-Gondwana derivation
and accreted to the Laurentia during the mid-Paleozoic (Rast and Skehan, 1983).
The Carolina superterrane is primarily composed of Neoproterozoic mafic to felsic volcanics of
volcanic arc origin and some volcaniclastic sedimentary rocks. Several plutons ranging in age
between 550 to 600 Ma intrude the arc complex (Hibbard et al., 2003). This entire assemblage
was metamorphosed to upper amphibolite to greenschist facies during Cambrian. This is
evidenced by ~ 530 Ma plutons that intrude the metamorphic rocks (Dennis and Wright, 1997).
Cambrian and Ordovician (identified from fossil assemblages) clastic sedimentary rock overlies
the arc assemblage (Samson, et al., 1982; Koeppen et al., 1995). The western Carolina
3
superterrane is metamorphosed to a higher grade with a 350-360 Ma metamorphic overprint and
contains numerous Devonian to Carboniferous younger granitoids and gabbros (McSween et al.,
1991; Hibbard et al., 2003) probably intruded during the docking of the Carolina terrane with the
Laurentia during mid to late Paleozoic.
1.3 Study Area and General Description of Lithotectonic Units
The study area includes the Hillabee greenstone belt, Ashland-Wedowee belt, Mulberry rock
gneiss and the Pumpkinvine Creek Formation (PCF) of the southern Appalachians (Fig. 1.1A).
Fresh unaltered samples were collected using several criteria including quality of mapping and
documented tectonic or stratigraphic significance of individual units.
Hillabee Greenstone:
The Hillabee Greenstone (HG) belt structurally occurs on top of the lower greenschist facies
Talladega Group in the Talladega belt. The Talladega Belt represents the most outboard
Laurentian margin cover sequence and the Talladega Group is the frontal metamorphic thrust
sheet of the southernmost Appalachians of Alabama and western Georgia. Like the PCF it is also
a sequence of bimodal volcanics (the maximum thickness is ~2.6 Km at places ) with the mafic
(~75%) and felsic members (~25%) being tholeiitic metabasalt and calcalkaline
metadacite/rhyolite respectively. The metavolcanic complex is in a thrust contact with the
underlying fossiliferous shallow marine Devonian to earliest Mississipian (?) metasedimentary
rocks of the Talladega belt.
The mafic member is primarily low-potassium tholeiitic metabasalt and basaltic meta-andesite
referred to as greenstone herein and the felsic member is metadacite referred as Hillabee dacite
herein (Tull et al., 2006). Unlike PCF there is not much of metasedimentary units present in the
HG belt. Rare, thin layers of micaceous quartzite and sericite phyllite occur locally within the
metavolcanic complex and are the only non-volcanic member of the HG belt.
Textural evidences like rarely preserved ophitic, intergranular, and porphyritic textures is
indicative of the extrusive nature of the mafic rocks in form of basalt flows and/or basaltic ash
(Tull and Stow, 1979, 1980; Stow, 1982). The felsic volcanics that are interlayered with the
mafic rocks occur in form of thick (upto 165 m) mineralogically and chemically homogeneous,
laterally continuous, tabular sheet-like bodies. Compositionally they are porphyritic meta
dacite/rhyolite. The metadacite shows bands of polycrystalline quartz alternating with
finegrained mica and opaque minerals. The rock contains porphyroclasts of hornblende,
actinolite, albitic plagioclase, and quartz (Tull et al., 1998) interpreted to be phenocrysts. Strain
is evidenced by undulose extinction of quartz grains and weakly kinked lamelleae of plagioclase.
Hillabee dacite extended over an area greater than 100's of sq. km. and are interpreted as largevolume pyroclastic ash flow crystal tuffs (Tull et al., 1998).
Samples of the greenstone and Hillabee dacite were collected for chemical analysis.
Structurally overlying the Talladega belt is the EBR allochthon.
4
Pumpkinvine Creek Formation: The Pumpkinvine Creek Formation is a linear belt of metavolcanic rock sequence extending from near the Alabama-Georgia border into northern Georgia
for a length of about 100 Km. It is included as part of >220 Km long Dahlonega Gold Belt
(DGB) terrane that is strike continuous and in the same structural position with the AshlandWedowee belt (Fig. 1.1, 1.1A). The importance of the PCF lies with the fact that it is structurally
adjacent to the Laurentian portion of the Appalachian making it the most inboard ‘suspect’
terrane (because of uncertain affinity, lack of basement it is considered to be “suspect” with
respect to Laurentia) in this portion of the southern Appalachians.
Lithologically the PCF is a sequence of bimodal volcanics in the outcrop scale and mapscale
containing fine grained amphibolite (referred as PCF amphibolites herein), felsic gneiss (termed
as Galts Ferry Gneiss or GFG member) and minor amounts of pelitic schists (Canton Schist) and
subordinate metamorphosed sandstone. The mafic and the felsic units are interlayered from
centimeter scale to tens of meters. The Canton schist possibly represent background
sedimentation (Holm, 2006) interpreted from its sporadic occurrence and in sharp contact with
the metavolcanic rocks of the PCF. McConnell (1980) identified discontinuous but regionally
mappable banded iron formation interlayered with the PCF amphibolites. Due to lack of any
ordered stratigraphic occurrence of the three units of the PCF (mafic, felsic and the
metasediments) they are all considered different facies of the PCF.
The amphibolite is fine grained, the major mineral phases being amphibole (75-80%) and
plagioclase (~20%) and the accessory mineral phases include epidote, garnet and actinolite
and/or chlorite (retrograde assemblage of amphibole). Typical volcanic textures include relic
amygdule filled with epidote and plagioclase phenocrysts are common. Pillow structures have
also been reported by several workers (McConnell, 1980; Abrams and McConnell, 1984).
The felsic metavolcanic lithology (GFG) occurs in two modes. In places it forms thin (0.1-0.5 m)
layers possibly representing rhyolitic ash eruptions. The second type of occurrence is as thick,
lenticular bodies up to 2 km thick extending along strike. Major mineral composition of the GFG
is plagioclase and quartz and enclaves of intermediate to mafic composition rich in amphibole is
common. Accessory minerals include muscovite and to a lesser degree garnet. The larger bodies
of felsic gneiss are likely large felsic eruptions.
The third and volumetrically least significant lithology of the PCF is the aluminous Canton
Schist, composed mainly of quartz sericite/muscovite ± garnet and biotite. This likely represents
background sedimentation during periods of volcanic quiescence (Holm, 2006). Presence of
extremely low sediments in archetypal back-arc basins such as the Lau basin-Havre trough
(portion of the Tonga-to-New Zealand back-arc system) have been reported by Gamble and
Wright, 1995.
In the northern portion of the study area the Chattahoochee fault (a late structure in the kinematic
sequence) demarcates the upper boundary of the PCF and emplaces higher grade, migmatitic,
predominantly metasedimentary units of the Sandy Springs Group (Higgins and McConnell,
1978) structurally above the PCF. The units structurally above the Pumpkinvine Creek
Formation, described as the New Georgia Group (Abrams and McConnell, 1981), contain a
mixture of pelitic and chemical sediments (iron-bearing quartzite), minor amphibolite and
5
locally altered ultramafic pods. These units also hosts numerous intrusive bodies varying in
composition from gabbroic to intermediate and felsic gneisses (eg. ~430 Ma Austell gneiss dated
by Higgins et al. 1997) and many more unnamed intrusive bodies. These units likely correlate to
the southwest into Alabama with part of the Ashland-Wedowee belt (Fig. 1.1). The intrusive
units are similar in age and composition to the granitoids (including the 460 Ma Kowaliga and
Zana granites) of Alabama as described by Russell et al. (1987) and Drummond et al. (1997).
The age of the New Georgia Group and correlative units of the Ashland-Wedowee belt to the
southwest is likely late-Proterozoic to earliest Paleozoic, and may represent slope-rise sediments
along the eastern margin of Laurentia (Tull ,1978). The intrusives might have formed because of
subduction-related slab melting beneath Laurentia (ex. Elkahatchee Quartz Diorite at 490 Ma) or
later magmatic events during Taconic and Acadian orogenies (Drummond et al.,1997).
Samples were collected from both the felsic (GFG) and mafic units of the PCF and Canton
Schist.
Ashland Wedowee Belt: East of the Talladega belt and west of the Brevard fault zone in
Alabama and west Georgia is a broad belt of medium to upper amphibolite facies schists,
gneisses and amphibolites of the Ashland – Wedowee belt. These rocks are interpreted to be late
Precambrian metamorphosed sedimentary rocks and minor basalts (Thomas et al., 1980). The
Ashland does not continue beyond the state line but the Wedowee and Emuckfaw extend
northeast into Georgia (Fig. 1.1A)
Exposed over much of the northern Alabama and western Georgia Piedmont (Fig 1.1, 1.1A) is a
distinctive graphite bearing metasedimentary sequence of slate, phyllite, quartzite and schist,
called the Wedowee Group, interpreted as the regressive phase of a more complex sedimentary
cycle involving many other rock units (Neathery, 1973). Protoliths for the metamorphic
lithologies possibly are arenite, siltstone and claystone interpreted to be a deep water
environment deposit.
The Ashland Super Group (Tull, 1978) underlies the Wedowee Group and is characterized by
metabasalts called the Poe Bridge Mountain amphibolites mostly concentrated in the lower part
of the sequence and metasediments with an aggregate thickness of few thousand meters. The
lower Ashland can be separated from the other units with relative ease because it is the only
stratigraphic unit containing abundant metabasalts (amphibolites). Ashland Super Group is a
thick sequence of metasedimentary rocks. Garnetiferous biotite schist and siliceous graphitic
muscovite schist are common. Less abundant rock types include kyanite-quartz schist,
sillimanite-quartz schist and quartzite (Vernon, 1973).
Although arial separation and apparent lithologic differences prohibit any direct correlation,
McConnell and Abrams (1984) speculated that the rocks of the new Georgia are at least in part
equivalent to the Ashland Supergroup. This is based primarily on the fact that both the New
Georgia Group and the Ashland Supergroup contain metavolcanic rocks and similar types of
sulfide hosted ore deposits. The authors also suggested that the rocks defined as Wedowee
formation in Alabama (Tull, 1978) are equivalent to the rocks of the Sandy Springs Group. This
correlation is based on lithologic similarities and the association of both Sandy Springs Group
6
and Wedowee Group with volcanic bearing rock groups (i.e., New Georgia Group and Ashland
Supergroup, respectively)
Fresh samples of quartzite, muscovite schist, garnetiferrous biotite schist from the Ashland and
Wedowee Group were analyzed for Nd isotopic composition to determine the model age of the
provenance.
Mulberry Rock Gneiss: Holm (2001, 2001a) studied the Mulberry Rock Gneiss (MRG), that
occurs as a structural window (eyelid window). The MRG occurs within a recess where the
eastern Blue Ridge rocks and ‘terrane bounding Allatoona fault are tightly folded into the
structural recess’ (Holm, 2001, 2001a). The unique position of the MRG (Fig. 1.1A) caused
previous workers to conclude that the yet undated rocks were Grenville basement to the
Talladega belt and the Talladega Group lies unconformably on the basement. Higgins et al., 1988
described them as equivalent to the nearby Grenville aged Corbin metagranite. Holm 2001,
2001a described MRG as medium grained, slightly metaluminous, two mica granite where
muscovite percent dominates over biotite. The other major minerals are quartz, plagioclase and
K-feldspar. Zircon, epidote and allanite occus as accessory phases. Tectono-stratigraphic models
of the MRG structural recess are tested in the light of the Rb-Sr whole rock age, U/Pb zircon
dates and Nd model ages of the MRG. The theories that are tested are either 1) whether MRG
represents Grenville basement to the Talladega Group or 2) it is one of the numerous Paleozoic
intrusion within the eastern Blue Ridge terrane.
Fresh samples of MRG was collected to look at the U-Pb age of the zircon, Rb-Sr whole rock
age and Nd model age to check the emplacement age as well as its connection with the Grenville
basement.
1.4 Purpose of the Study
This study will focus on the geochemical and geochronological constraints on the origin and
evolution of different tectonic divisions of the eastern Blue Ridge, southern Appalachians. A
major problem in Appalachian geology is the lack of Taconian deformation in the southern
Appalachians. The eastern margin of Laurentia was curved into bays and promontories during
opening and closing of the Iapetus ocean. Talladega belt (now within Alabama recess)
originated as a continental promontory. During continental collision, the protruding promontories
should first record the deformation. The absence of any Taconic deformation events in the
Talladega belt rocks (outermost preserved portions of the Laurentian margin in the southern
Appalachian) implies that a major promontory of the Laurentian margin might have escaped the
effects of Taconic orogeny (Tull, 1998). As in the central Appalachians, it is difficult to appeal to
subduction beneath a proven Ordovician arc as a deformation mechanism for the southern
Appalachians. This study will use geochemical and geochronological data supported by detailed
field maps to investigate whether or not the PCF formation can qualify for the Ordovician Arc.
The PCF and Hillabee, exist only as fragments, unlike the Taconic arcs of the New England
Appalachians (e.g. Bronson Hill arc) and the Ordovician northern Appalachian arcs (e.g.
Chopawamsic Terrane) that commonly include typical arc related features such as voluminous
arc-related sediments, accretionary wedge, and evidence of accretion in the foreland. If the PCF
and Hillabee are the remnants of the Ordovician volcanic arc, the tectonic setting, timing and
7
style of the Taconic orogeny in southern Appalachians and formation of the arc-related terranes
took place in an environment very different from the Ordovician arcs of the northern and central
Appalachians. Possibly the arc in the southern Appalachian formed during Ordovician but
remained outboard of the Laurentian margin until terminal closure of the Iapetan Ocean.
The goal of this research is to propose a tectonic model for the southern Appalachians during
Taconic orogeny. The study will be based on new detrital/magmatic zircon ages, major and trace
element analysis and Sr and Nd isotopic compositions of major tectonic units from southern
Appalachians tectonic divisions west of the Brevard fault zone.
1.5 Analytical Techniques
1.5.1 ICP MS Trace Element Analysis
Trace elements for the PCF amphibolites, HG, GFG and Hillabee metadacite were determined by
solution ICP-MS analysis using a ThermoFinnigan Element. Small chips of each sample were
handpicked and crushed in agate mortar and pestle. Thirty mg of each sample was weighed into
screw top Teflon beakers and dissolved with 2 ml of 3:1 distilled HF-HNO3. After drying on a
hot plate at 120oC, the samples were allowed to reflux overnight in concentrated HNO3. Samples
were then dried again and brought to a final volume of 60 ml with 2% HNO3. The sample
solutions were further diluted in the ICP vials for target concentration of 100 ppm TDS. One ppb
of Indium was added as internal standards and drift corrections for each analyzed mass were
applied by interpolating with the internal standard. The solution ICPMS analyses of the
amphibolites and greenstones were calibrated against a single solution of well-characterized
Hawaiian basalt USGS standard BCR-1 prepared identically to the samples. For the felsic
samples G2 was used as the USGS standard. BHVO-1 was also prepared like a sample and
calibrated against the BCR-1 and G2 to check the precession of measurement.
1.5.2 Sr-Nd Isotopic Analysis
Fresh whole rock samples chosen for whole rock analysis of Sr and Nd were powdered in an
agate mortar and dissolved in a 3:1 mixture of 2X distilled HF: HNO3. Separation of Sr and Nd
followed ion exchange procedures employing 4.5 ml of AG50W-X8 9-cm bed-length ion
exchange resin. Nd was separated as a bulk REE fraction and eluted in 6 N HCl. Nd and the REE
fraction was further separated on a 1.2 ml, 6 cm bed-length column of Ln resin SPS.
Measurements were made on a Finnigan-MAT 262 mass spectrometer. Strontium measurements
are reported relative to the measured value of the E&A Sr standard of 87Sr/86Sr = 0.708000 ± 11
(2sd, n = 15) and Nd relative to the measured LaJolla standard 143Nd/144Nd = 0.511848 ± 11
(2sd, n = 12). n= number of analysis, each analysis being the average of 100 ratios.
1.5.3 Rb/Sr Ratio Measurement in ICP MS
A precise method for Rb and Sr ratio measurements for samples (GFG) with low Rb/Sr ratios
(<10) has been developed. Fresh whole rock samples were powdered in a agate mortar and
8
dissolved in a 3:1 mixture of 2X distilled HF: HNO3 and then diluted to obtain Rb and Sr
concentrations in the 10 ng/ml range. These solutions were then directly measured without
elemental separation of Rb and Sr from the rock matrix or from one another. Measurements of
the peaks: 85Rb+, 86Sr+, and 88Sr+, were performed by ICP-MS, using a ThermoFinnigan Element
in low-resolution mode. Ion count rates were in the 10-100 MHz range. Isobaric interference of
86
Kr+ from the Ar plasma gas was monitored on acid blank solutions, and was found to be
negligible (<100 Hz). The precision on standard solutions was ±0.2%. The method was applied
to two USGS rock standards, G-2 and AGV-1, and yielded Rb/Sr ratios that were accurate to
within 2-4%. These data are within the reported uncertainties of the Rb/Sr ratios of the two
standards (Gladney et al., 1988).
1.5.4 Rb/Sr Ratio Measurement by Isotope Dilution
The MRG samples have a higher range of Rb/Sr ratios and were measured by standard ID
technique. Approximately 100 mg of powder was spiked with isotopically enriched tracers of
84
Sr and 87Rb. The samples were then dissolved in screw cap teflon beakers with distilled 3:1 HF
and Nitric acid. The dissolved samples were converted into a chloride form and Rb and Sr were
separated using a cation exchange column containing AG 50W-X8 resin with a volume elution
of 2.5N HCl. The total analytical blanks for Sr and Rb were less than 100 picogram and 1
nannogram respectively.
Isotopic analysis were made on a Finnigan-MAT 262 mass spectrometer. Rb and Sr were loaded
on W- single filament with TAPH (10g H2O + 0.5g TaCl5 + 3g 0.1N H3PO4 + 0.5g conc. HF)
solution. The measured Sr isotopic compositions were corrected for the isotopic composition of
the spike and normalized to a 87Sr/86Sr value of 0.1194.
1.5.5 LA-ICP-MS Single Zircon Analysis
Samples were collected from GFG, MRG and a thin meta-sandstone unit interlayered with the
Canton Schist. At each sample locality, 1–2 kg of fresh whole rock samples were collected.
Samples were prepared for analysis using standard crushing and separation techniques, including
heavy liquids and magnetic separation. Apparently inclusion-free zircons were then hand-picked
under a binocular microscope. At least 100 zircons from each sample were mounted in epoxy
and polished.
U–Pb analyses were performed at University of Arizona, Tucson with a Micromass Isoprobe
multicollector Inductively Coupled Plasma Mass Spectrometer (ICP-MS) equipped with nine
Faraday collectors, an axial Daly collector, and four ion-counting channels. Zircons were ablated
with a New Wave DUV 193 nm Excimer laser ablation system. All analyses were conducted in
static mode with a laser beam diameter ranging from 30 to 50 μm. Contribution of Hg to the
204
Pb mass position was removed by subtracting measured background values. Isotopic
fractionation was monitored by analyzing a Srilankan zircon standard, which has a concordant
TIMS age of 564 ± 4 Ma (Dickinson and Gehrels, 2003). This standard was analyzed once for
every five unknowns in detrital grains and once for every three unknowns in magmatic zircons.
Uranium and Thorium concentrations were monitored by analyzing a standard (NIST 610 Glass)
with approximately 500 ppm Th and U. The calibration correction used for the analyses was 2–
9
3% for 206Pb / 238U and approximately 2% for 206Pb / 207Pb (2-sigma errors). The lead isotopic
ratios were corrected for common Pb, using the measured 204Pb, assuming an initial Pb
composition according to Stacey and Kramers (1975) and uncertainties of 1.0, 0.3 and 2.0 for
206
Pb / 204Pb, 207Pb / 204Pb, and 208Pb / 204Pb, respectively.
For samples that are younger than 1 Ga ages are considered reliable if five or more analyses
performed in different grains yield overlapping 206Pb / 238U ages. This strategy is used because of
the low precision of 206Pb / 207Pb ages for young grains, making concordance/discordance a poor
criteria for determining reliability. Clustering is also a better criteria for reliability than
concordance given that Pb loss and inheritance in young systems can create concordant ages that
are significantly younger or older than the true ages. Such analyses could be concordant but
would not define a cluster, and would accordingly be rejected as unreliable.
Three samples were analyzed for U-Pb geochronology in this study, GFG, MRG and detrital
zircon from a meta-sandstone. Ages of 50-60 zircon grains were measured from each igneous
sample and ~100 grains were measured from the sedimentary sample. Results and errors are
reported in Tables 1.7, 1.8 and 1. 9. Each line in the Table represents a spot analysis.
1.6 Results
1.6.1 Major and trace-element geochemistry
The major and trace element concentrations of the PCF, greenstone, GFG and Hillabee dacite are
shown in Tables 1.1, 1.2 and 1.3.
The bimodal nature in terms of SiO2 of both mafic and felsic components of the PCF and the HG
sequence is shown in a histogram (Fig.1.2). It is important to note that there is no intermediate
component associated with these units. Using the total alkalis vs. SiO2 diagram (Fig.1.3) by
Lebas et al. (1986), the PCF amphibolites and greenstone samples plot within the basalt field and
the GFG and Hillabee dacite plots in the rhyolite and dacite field respectively. Tull et al., (1978)
reported some basaltic andesite from the Hillabee that are plotted in the diagram too. The GFG
composition is extremely sodic (Table 1.1) and is strongly peraluminous (Fig. 1.4A). The
Hillabee dacites also plot in the peraluminous field but they are more near the boundary between
metaluminous and peraluminous field. Traditionally it is thought that strongly peraluminous
magmas are derived from mature pelitic sediments and produce S-type igneous rocks. Miller
(1985) describes more likely sources of these rock types to be continentally derived immature
sediments and intermediate to felsic igneous rocks as well as certain products of partial melting
in metaluminous mafic sources. Miller (1985) describes criteria for identifying S-type granites
(Chappell and White, 1974) and pelitic parentage of igneous rocks. The GFG and Hillabee
metadacite do not fit either criteria and when compared to a typical pelitic composition (Gromet
et al., 1984), show depletion in rare earth element (REE) chemistry.
In the AFM diagram (Fig. 1.4B) the amphibolites of both the Hillabee and PCF defines a
tholeiitic trend where as the felsic components follow the calc alkaline trend. In the Ti vs. Zr
plot (Fig. 1.5A & 1.5B) the felsic components and the amphibolite shows two parallel trends
depicting some amount of differentiation within each group but indicates that the felsic
10
component is possibly not derived by simple fractional crystallization of the associated mafics.
Further discrimination of the PCF and greenstone using trace element concentrations will help
elucidate its origin.
Commonly used geochemical parameters in distinguishing arc and marginal basin rocks include
light rare earth element (LREE) fractionation leading to enrichment and Nb-Ta anomalies
relative to MORB values. High field strength elements (HFSE) provide values typically
unmodified by moderate grade metamorphism. These elements are used as robust parameters in
discrimination of tectonic environment (Cabanis and Lecolle, 1989; Meschede, 1986; Pearce and
Norry, 1979; Pearce et al., 1984; Shervais, 1982; Wood 1980).
1.6.1a PCF Amphibolites and Greenstones.
Because the rocks of the PCF have undergone amphibolite grade metamorphism and the
greenstone belongs to the lower green schist facies, it is essential to check the mobility of
elements during regional metamorphism and subsequent alteration. Using covariation diagrams
(Kim et al., 2003) plotting significant elements against HFSE such as Zr, a linear trend suggests
relative immobility of those elements. Figure 1.6 illustrates linear trends between Nb, Y, V, and
the HFSE Zr in PCF and Hillabee mafic facies, thus signifying immobility of the elements used
for tectonic discrimination and their utility in identifying tectonic environment of emplacement.
Several discrimination diagrams have been developed for mafic rocks based on HFSE including
Wood (1980), Pearce et al. (1984), Meschede (1986), Cabanis and Lecolle (1989), and Shervais
(1982). Discrimination diagrams (Figs. 1.7A,B) using HFSE including Meschede (1986) and
Pearce et al. (1984) help constrain the PCF amphibolites and greenstone as island-arc related
yielding compositions in the range of N-MORB to volcanic arc basalt. Cabanis and Lecolle
(1989) use the relationship between La and Nb to help discriminate between volcanic arc
signatures and N-MORB signatures and include a field for BABB (Fig. 1.8) where PCF
amphibolite and greenstone samples fall. High La relative to MORB is a signature of SSZ
compositions (Hawkins, 1995, 2003).
Decoupling of large ion lithophile elements (LILE) and HFSE in the supra subduction zone
environment, termed “flux melting” (Gill, 1981), yields enrichment of LILE such as Ba, Rb, and
Sr in BABB and is likely due to dehydration fluids associated with the subducting slab while
HFSE or HREE remain unaffected because of their immobility in aqueous fluids (Hawkins,
1995, 2003). Flux melting paired with decompression melting melts more of the hydrous mantle
wedge than either process alone can generate (Xia et al., 2003). When normalized to Primitive
Mantle, PCF amphibolites display a modest LILE enrichment and HFSE near Primitive Mantle
values (Fig. 1.10). Another typical BABB characteristic is the negative Ta-Nb anomaly (Kim et
al., 2003). Rare earth element (REE) patterns of the PCF are extremely flat exhibiting minimal
fractionation, which is typical of tholeiitic basalt (Fig. 1.11). The REE’s are considered relatively
immobile and would not yield a supra subduction zone signature.
The greenstone samples shows enrichment in LILE relative to the Primitive Mantle and a Nb
depletion but the primary difference with the PCF samples being Pb enrichment which is one of
the important features of the arc rocks (McCulloch and Gamble, 1991) (Fig 1.10). 3 out of 7
samples show moderate Zr and Ti anomaly, another characteristics of arc basalts. The REE
11
pattern are flat, similar to the PCF amphibolites though they show higher range of variability
possibly due to olivine fractionation at the source (Fig. 1.11). Due to geochemical similarities
and similar tectonic settings the PCF and Hillabee volcanic sequence might represent different
parts of the same arc-back arc system.
1.6.1b Galts Ferry Gneiss and Hillabee dacite.
Maniar and Piccoli (1989) recommend the use of major element data in tectonic discrimination
of granitoids, however, elemental mobility in multiply deformed and metamorphosed terranes in
the southern Appalachians limits this approach. In an analogous approach to that of basaltic
rocks, Pearce et al. (1984) use multi-element plots emphasizing HFSE as parameters and
normalize samples to hypothetical ocean ridge granite (ORG) in order to discriminate tectonic
fields. In Nb versus Y and Rb versus Y + Nb (Fig. 1.9) and Yb versus Ta and Yb + Ta versus Rb
(not shown), the GFG and the Hillabee metadacite plots in the volcanic arc granite field. All
GFG samples exhibit moderate light rare earth element (LREE) enrichment and a modest Eu
anomaly (Fig. 1.12) except for sample SC318G that is a very thin (<0.1 m) felsic layer and was
probably an ash deposit. A couple of samples show slight positive Eu anomaly that might be due
to the presence of plagioclase phenocrysts.
The Hillabee dacites are LREE enriched with (Ce/Yb)n ranging from 11.2 to 14.3. The resulting
REE pattern thus display a distinct negative LREE slope typical of calc-alkaline dacites in
orogenic belts (e.g. Gill,1976). Absence of Eu anomaly indicate that plagioclase fractionation is
not the cause of geochemical variations.
1.6.2 Sr and Nd Model Values
The Nd and Sr isotope systems are most useful in providing insight as to the nature of, and
mixing between, sources during magmagenesis. The Sr and Nd values of the PCF amphibolites,
GFG, greenstone and Hillabee dacite are shown in Table 1.4.
The measured Sr and Nd isotopes are recalculated at 460 Ma (Russell, 1978; Thomas et al,
2001). The PCF amphibolite samples exhibit epsilon Nd (T460 Ma) values of +3.3 to +7.7 and the
greenstone exhibit εNd (T460 Ma) values ranging between +2.01 to +6.84, typical of a depleted
mantle source composition (Table 1.4). The GFG epsilon Nd (T460 Ma) values of –3.21 to 4.65
and the even more negative Hillabee dacites (ε Nd at T460 Ma between –3.9 to –5.5) indicate that
this magma was not derived directly from fractional crystallization of the magma that formed the
amphibolite protolith. It is likely evidence of sampling of a more isotopically evolved component
(Fig. 1.13). This may include supra subduction zone fluids, subducted sediments, assimilation of
the older basement rocks or melting of the basement.
1.6.3 Rb-Sr Whole Rock Age
Rb-Sr whole rock analyses of the GFG (Table 1.5) are employed to help constrain timing of
crystallization and possibly provide insight into petrogenesis of the felsic facies of the PCF. Fig.
1.14 is a plot of the 87Sr/86Sr vs. 87Rb/86Sr data. For reference a hypothetical 391 Ma isochron is
drawn. Out of the eight samples analyzed, they appear to cluster into two groups with different
initial 87Sr/86Sr ratios. The two groups can be separated based on their REE pattern too (Fig.
12
1.15). Four samples with low 87Sr/86Sr initial (~0.705), similar to the associated amphibolites
have a relatively flat LREE and HREE pattern (La/Yb = 1.2 to 4.2) where as the samples with
higher initial 87Sr/86Sr (~0.708) displays a steeper slope for the LREE (La/Yb = 4.2 to 23.7). It
is likely that the magmagenesis is complex and may tap into multiple sources. The GFG samples
possibly define two isochrons of similar ages (424±25 Ma and 450±66 Ma) with significant
errors but differing initial 87Sr/86Sr. The age within the error margin might represent either the
crystallization age or the regional metamorphic age of 366 ± 25 Ma (Russell, 1978) of the EBR
metasediments. A lower 87Sr/86Sr initial value of 0.7059 is similar to the associated (PCF) metabasalts and is likely closely related to the primary melt by fractional crystallization of the metabasalts and subsequent eruption (Fig. 1.14). A higher 87Sr/86Sr initial value of 0.7082 proves the
existence of another, more radiogenic/isotopically-evolved, source bearing differing initial ratios
that may include a mixed (mantle source plus recycled crust and/or sediment) source, possibly by
assimilation and fractional crystallization (AFC). There are also the possibilities of Sr mobility
due to subsequent thermo-tectonic events, or initial 87Sr/86Sr variability in the magma source.
The Hillabee metadacites yielded a Rb-Sr whole rock age of 395 ± 20 Ma with a initial 87Sr/86Sr
ratio of 0.7082 (Durham, 1993). This age is very close to K-Ar whole rock age of 399 ± 17 Ma
of the slates of the Talladega Slate Belt reported by Kish, 1990 and within the range of
uncertainty for the whole rock Rb-Sr age of 366 ± 25 Ma for the Ashland-Wedowee
metasediments and the K-Ar age of 382 ± 14 Ma for hornblende in the Hillabee metadacite
reported by Russell, 1978. Thus Durham interpreted the Rb-Sr whole rock age of the metadacites
as the age of lower green schist facies regional metamorphism. The initial 87Sr/86Sr ratio of
0.7082 is similar to the higher initial values of the GFG. The high initial Sr isotopic ratio
observed for the GFG and Hillabee metadacite is within the range of ratios observed by Pushkar
et al., 1973 in the Lesser Antilles. The higher ratios (0.709) observed in the Lesser Antilles was
explained by assimilation of crustal material, either continental detritus or marine sediments.
1.6.4 Zircon Data
Incorporation of older crustal material by the PCF volcanics can be tested by looking at the
zircon inheritance (if any) of the GFG (Table 1.7). Zircons from the GFG samples were
analyzed optically under a petrographic microscope and a scanning electron microscope (SEM)
in backscattered electron (BSE) mode. Approximately 50 zircons were analyzed and almost all
zircons yielded Ordovician ages and displayed igneous morphologies (e.g. euheadral crystals)
with no detectable optical zoning. When plotted in a histogram the modal class ranged from 440
to 480 Ma (206Pb/238U age) that comprised of ~ 35 grains (Fig 1.16). A weighted 206Pb/238U
average of the 35 grains yielded an age of 460.2 ± 9.7 Ma. This agrees with the 463 ± 3 Ma age
of the GFG (near Allatoona Dam) reported by Thomas et al., (2001). A few grains (n~4) under
the Cathode Luminescence (CL) image show the presence of cores in the zircons (Fig. 1.16A).
The core age is Cambrian. No older zircon (Grenville age) was found. Zircons from Hillabee
metadacites were not analyzed in this study but Russel, 1978; Russel et al., 1984, McClellan and
Miller (2000) and Tull et al., (2006) reported an Ordovician age (~460-470 Ma) for the complex.
Some of the ion probe data set on the Hillabee metadacite (McClellan et al., 2005) shows older
Cambrian core at ~500 Ma., i.e. 30-40 m.y. older than the interpreted Hillabee crystallization age
similar to that observed for the GFG.
13
1.7 Nd Model Age and Detrital Zircon Ages of the Metasediments of the
eastern Blue Ridge.
A direct approach of testing the sediment source is looking at their Nd-model age. Nd and Sm
may undergo appreciable fractionation during extraction of crustal material from the mantle but
are typically not fractionated during process such as erosion and metamorphism. Sm-Nd isotopic
data for a rock can be used to calculate a model age representing either a crustal extraction or
average crustal residence age. The significance of the model age depends on the proportion of
the juvenile verses ancient sources contributing to the rock. A TDM age is the time when a
sample and the mantle source had identical isotopic composition and a TCHUR represents the
extraction of the source rock from the Chondritic Uniform Reservoir (CHUR). In practice, Nd
isotopic data and Nd model ages are used as chemical finger prints to distinguish between crustal
provinces and terranes within an orogen (e.g. Patchett & Ruiz, 1989; Bennett & DePaolo, 1987).
Table 1.6 summarizes isotopic data from the metasediments of the Ashland-Wedowee belt and
the Canton Schist that is interlayered with the PCF. The Ashland-Wedowee sediments generally
exhibit a range in TCHUR ages from 964 m.y to 1119 Ma with one sample at 1439 Ma. Canton
Schists range in age between 943 to 1137 Ma.
In order to make the Nd model age internally more consistent a depleted mantle (DM) extraction
model was tested. Nd isotopic values and Sm/Nd ratios of the DM for the range of ages for the
depletion event of the two-stage evolution model were taken from Salters and Stracke 2004.
Models were tested for a minimum depletion age of 1.8 Ga, which is similar to the minimum age
expected from the Pb-isotope systematics in MORB (Hart, 1984; Tatsumoto, 1978). A maximum
depletion age of 3.4 Ga inferred from the Re-Os isotope systematics. A 2.2 Ga was based on the
assumption that continental crust extraction occurred at 2.2 Ga (Chase and Patchett, 1988;
Condie, 2000; Galer et al., 1989) and if the DM is assumed to be complementary to the
continental crust a similar average age is expected. The 3.4 Ga depleted mantle formation fits the
model best (Fig. 1.17) and 12 out of 15 metasedimentary samples analyzed range in age between
1514 – 1580 Ma, two samples are slightly younger at 1450 Ma and 1482 Ma and one sample is
1892 Ma (Table 1.6).
The difference in the model age might be due to difference in source character. In order to test
the hypothesis the REE’s of the metasediments are plotted in a North American Shale
normalized diagram (Fig. 1.18, Table 1.10). The normalized REE patterns of the
metasedimentary samples are flat, do not show any visible difference and are 0.5 to 2.5 times
more enriched than the North American Shale with the exception of the two young samples that
have the highest concentrations of all the REEs. The sample with the oldest age has a flat REE
pattern similar to the North American Shale.
A suite of detrital zircons (n~100) separated from a meta-sandstone interlayered with the Canton
Schist were analyzed in the same method as the GFG samples (Table 1.8). A concordia diagram
and histogram derived from LA-MC-ICPMS ages of the detrital zircon grains shown in Fig 1.19.
The dominant feature of the detrital zircon age distribution is a pronounced cluster of ages
between 918 and 1190 which contains nearly eighty percent of the ages. Zircons of this age
range are clearly the dominant signature of the Grenville terrane, as they are the most abundant
age group. Age peaks at 511 Ma and 589 Ma (n=3) are somewhat curious as they are not ages
representative of Laurentian zircon populations. In Newfoundland, Cawood and Nemchin (2001)
14
report synrift (572-628 Ma) detrital zircon with a local igneous source. Older peaks of 1351 Ma
and 1462 Ma and also the TDM age of 1439 of one Ashland sample correspond to the ages of the
eastern Granite-Rhyolite province (Becker at al., 2005).
1.8 Mulberry Rock Gneiss
1.8.1 Rb-Sr Results
The Rb-Sr whole rock isochron for MRG is shown in Figure 1.20. The analytical error is smaller
than the sizes of individual data points. The calculated isochron age is 467 Ma with an
uncertainty of 16 Ma at 95% confidence level.
The range of 87Rb/86Sr, from 5.1 to 51.9 is very large in this body (Table 1. 5). This suggests that
any metamorphic events that led to the current fabric of the gneiss did not chemically or
isotopically homogenize the body, at least not sufficiently to produce an isochron reflecting the
time of metamorphism. More likely the linear isochron reflects the time of crystallization of the
MRG from a granitic magma. The initial 87Sr/86Sr is 0.7076±0.0036. This value suggests that the
source of the MRG magmas involve some fraction of crustal rocks.
1.8.2 Nd Model Ages.
Table 1.4 summarizes the Nd-isotope data that exhibit a range in εNd of the MRG between –1.6
to –2.8. Thus it can be interpreted that the MRG had recycled older (Laurentian?) crust.
1.8.3 U-Pb Zircon Ages.
Zircon grains (n~70) were separated from the MRG samples and analyzed by LA-MC-ICPMS.
The 206Pb/238U ages show quiet a range with the highest number of grains plotting in the age
range between 420 and 450 Ma (histogram plot of Figure 1.21, Table 1.9). Inherited
components (?) and/or partially re-set older zircon xenocrysts have a range of ages from 450 to
600 Ma, with one concordant grain at 1068 ±20 Ma (Pb-Pb age-See Table). Pb loss seems to be a
problem too for these samples with 30% of the grains ranging in age between 300 and 400 Ma.
On a concordia diagram the error ellipses cluster around ~450 Ma with the upper intercept of the
discordia line at 1049 Ma and the lower intercept at 345 Ma. The upper intercept represents
Grenville inheritance while the lower intercept might be lead loss due to a thermal event. Thus in
light of the Rb-Sr whole rock data, Nd isotope data and U-Pb zircon analysis it can be concluded
that MRG is not part of the Grenville basement, however it might have originated by anatectic
melting of older Grenville-aged rocks.
1.9 Discussion
1.9.1 Petrogenesis of Modern and Ancient Arc/Back-Arc Systems – Origin of the Bimodal
Volcanics
15
It has been expressed that arc and back-arc basins appear to be simple tectonic environments, yet
they often display tectonic and geochemical heterogeneity on the scale of kilometers and
sometimes even within a single volcanic center (Gamble and Wright, 1995). Behind the arc, a
back-arc basin is typically developed. This is the site of MORB-like volcanism that creates thin
ocean type crust in an extentional tectonic environment behind the volcanic arc. It has been
suggested that felsic volcanism accompanies early rifting in the back-arc setting (Marsaglia,
1995) yielding bimodal volcanism with an arc-like signature during these early stages, whereas
the more mature back-arc basin volcanism displays mid-ocean ridge basalt (MORB)-like
characteristics (Xia et al. 2003). In Xia et al.’s (2003) model of an Ordovician back-arc basin in
the northern Qilian Mountains of China, it is suggested that early back arc rifting does not
influence normal arc magmagenesis in the mantle. During initial rifting, some of the arc melts
are diverted to the young rift axis yielding the possibility of incorporation of supra-subduction
zone (SSZ) components and melting of the ambient mantle. A substantial backarc extension is
necessary before induced mantle upwelling is sufficient to generate backarc-spreading-ridge
lavas (White and McKenzie, 1989). From this perspective, it is impossible to generate backarcspreading-ridge lavas during the earliest island-arc rifting stages of back-arc-basin formation.
Continued back-arc spreading and the consequential relative movement away from the locus of
arc magmagenesis should yield normal mid-ocean ridge basalt (N-MORB) -like composition of
the newly formed crust. Descriptive terminology of back-arc basin basalts (BABB) is variable in
that it includes an array of arc-like to MORB-like compositions. Taylor (1995) suggests that
modern back-arc volcanism exhibits diversity in the generation of magmas at back-arc spreading
ridges and feels that it is not feasible to describe compositional differences in basalts relating to
stages of maturity of a back-arc basin. There are probably three first-order components that
influence the rock chemistry of evolving back–arc basins including melting of the incipient
island arc during initial rifting, supra subduction zone fluids associated with the dehydration
melting of the subducting oceanic slab and decompression melting associated with the mantle
wedge. The first two components should be most important during the initial and early back-arc
spreading and should decrease over back-arc maturation as the spreading axis progresses away
from the arc and SSZ fluid source. Many back-arc basins, ancient and modern, show variation
along the axis of back-arc spreading from the incipient rifting, yielding island arc-like basalts, to
a mature back-arc where basalts are MORB-like (Xia et al., 2003; Gamble and Wright, 1995;
Hawkins, 1995 and 1976) giving rise to the term back-arc basin basalt (BABB). It has also been
shown that MORB-like basalts can occur concomitantly and proximal to arc-like eruptions in the
initial spreading of an active back-arc (Fryer, 1992).
Modern or young analogues of back-arc bimodal magmatism are limited. Cole et al. (1995)
describe the back-arc volcanism in the Taupo Volcanic Zone, located at the North Island of New
Zealand as the southwest extension of the oceanic back-arc Havre Trough where magmagenesis
of a large volume of rhyolite here is likely composed of crustal materials incorporated with a
mantle source. The Okinawa Trough exhibits a bimodal petrogenetic nature. Shinjo and Kato
(2000) put forth a model explaining the origins of basaltic, rhyolitic and hybrid magmas
generated at the Okinawa Trough. The authors have identified two types of rhyolite (and
associated basalts) and hybrid andesite related to different petrogenetic processes. While the
Okinawa Trough is a continental back-arc basin, the modes of magmagenesis are limited to
fractional crystallization of the associated basalts, assimilation and fractional crystallization
(AFC) of basaltic magma, and hybridization of evolved basalts and rhyolite derived from
16
fractional crystallization. All these processes involve a lower gabbroic crust. Silurian bimodal
volcanism is described in the northern Appalachian Coastal volcanic belt of New England and
New Brunswick. While it is documented as a Silurian continental extensional environment (Van
Wagoner et al., 2002), identification as a back-arc basin remains enigmatic because of
continental within plate geochemical signatures and island arc-like Nb anomalies. It is common
for bimodal volcanism and the majority of highly silicic volcanism to occur in the early stages of
rifting in continental regimes by underplating and anatexis of pre-existing continental crust and
also in intra-oceanic regimes. Ringwood (1977) described the petrogenesis of a tholeiitic suite of
rocks concomitantly with a calc-alkaline suite that is derived from partial melting of the tholeiitic
magmas in source regions comprised of either subducted oceanic crust or mantle wedge
materials and yield an initially high silica content.
Subduction is a complex process that produces some characteristics igneous association with
distinctive pattern of the major and trace element chemistry. Andesites and basaltic andesites
dominate the island arc rock spectrum. Basalts, dacites and rhyolites are subordinate. In case of
the PCF and the Hillabee greenstone sequence mostly basalts, dacites and rhyolites are present.
Tull et al., (1978) reported some basaltic andesite in the Hillabee greenstone sequence. This
apparent discrepancy in the lithologic association could be explained by the deep erosion that the
PVC and the Hillabee sequence suffered and are preserved today as thin slivers of the remnant
arc. If the exposed plutons of the eastern Blue Ridge represents the core of the arc (Drummond et
al., 1997) then the entire arc carapace is eroded which might have been the andesitic and basaltic
andesitic members of the Ordovician arc.
The petrogenetic relationship between the felsic and mafic facies of the PCF and Hillabee is
important in understanding the arc/back-arc evolution. It is difficult to determine whether these
facies follow the same trend between tholeiitic and calc-alkaline because there are no
intermediate values between the suites. The mafic facies displays a tholeiitic trend whereas the
felsic GFG and the Hillabee dacite displays a calc-alkaline trend as shown in the AFM diagram
(Fig. 1.4B). Shinjo and Kato (2000) use mass balance models to identify the contribution of
fractional crystallization and assimilation/fractional crystallization (AFC) in the Okinawa Trough
bimodal back-arc volcanic suite. It is important to identify a parameter useful in discriminating
the process of generation of the felsic component. Rollinson and Windley (1980) use the
relationship of Zr and Ti as an indicator of fractional crystallization as a source mechanism of
felsic melt during magmagenesis. If fractional crystallization of felsic material from a mafic
source is the main contributor to the genesis of the felsic magmas, the felsic material would
either fall along the same trend-line as the mafic material or, after Ti-bearing minerals such as
magnetite or titanite crystallize out, the felsic material would fall off the trend but with
increasing Zr content. In figure 1.5A and 1.5B, it is clear that the felsic does not fall along the
differentiation by fractional crystallization trend of the spatially related mafic material. It is
unlikely that fractional crystallization is responsible for the generation of the felsic magma. It is
important to examine possible source mechanisms for the GFG and the dacites. As it is unlikely
that fractional crystallization is the major factor in genesis of the GFG, it is important to explore
the possibility of supra subduction zone (SSZ) components such as dehydration and partial
melting of a subducting slab, mantle wedge components and later decompression melting at a
spreading ridge or partial melting of arc crust or a preexisting older continental crust. Further
17
discrimination of the PCF and greenstone by major and trace element will help elucidate its
origin.
The major chemical difference between the two felsic members, the GFG and the Hillabee
dacites is in their silica content and alkali content. GFG has low content of potassium (K2O<2%)
and high Na2O making it a sodium rhyolite/ trondhjemite. The flat REE pattern of the GFG
((La/Sm)n ranges from 1.54 to 4.47) also indicates fractionation from a mafic source. The
Hillabee dacites have flat HREE and enriched LREE ((La/Sm)n ranges from 4.60 to 5.73) typical
of calc alkaline magmas. Thus the dacites might be generated by a different process than the
GFG which involved either partial melting of lower continental crust and/or interaction with the
Laurentian continental crust.
Because back-arc environments are complex and the compositions of magmas produced in these
environments are variable with back-arc maturity, it is difficult to distinguish the contributing
source to the GFG and Hillabee dacite magmas here. Because back-arc basins in their initial
stages often rift an existing arc, it is conceivable that the GFG and Hillabee dacite magma
includes supra subduction zone fluids, mantle wedge materials, and assimilation of the rifted arc.
The possible variability in the source material makes it difficult to complete a mass balance of
the felsic volcanics based on typical epsilon Nd values of these sources. Initial 87Sr/86Sr values of
the PCF amphibolite vary between 0.7040 and 0.7069, though mostly cluster around 0.7045. The
greenstone has a spread between 0.7039 to 0.7063 with a cluster around 0.7050. Variability
could easily be caused by Sr or Rb mobility during metamorphism. Initial 87Sr/86Sr of the GFG
samples are scattered between 0.7046 and 0.7084 while a low population, <0.706, and a high
population, >0.708, exist (Figure 1.13). The population distribution within the GFG can be
derived from Sr mobility during metamorphism, but could also further indicate that different
sources contributing to the magma. GFG samples bearing similar initial 87Sr/86Sr values (~0.705)
to that of PCF amphibolites might represent a liquid line of decent from PCF magma, whereas
initial 87Sr/86Sr values >0.708 are likely derived from a more isotopically evolved source. The
metadacites however has a smaller spread for initial 87Sr/86Sr ranging between 0.7069 to 0.7082
and distinctly crystallized from a more evolved source than the associated greenstones.
Many of the Ordovician arc-related rocks of the Appalachians show evidence of interaction with
Laurentian (Grenville-aged) continental crust. The abundance of inherited zircons and very
negative epsilon Nd values attest to the close interaction of these arcs with the preexisting
Grenville/Laurentian crust (Coler et al. 2000).
It is important to attempt to frame the possible relationships to other concomitant arc terranes in
order to determine a possible reference frame for arcs eventually accreted onto Laurentia. Most
Appalachian arc-related data is available from the northern portion of the orogen, and in a
compilation, by Coler et al. (2000) and references therein, of available Nd isotopic data, a Nd
isotopic evolution diagram (Fig. 1.22) shows that most northern Appalachian (and two southern
Appalachian) arcs and arc-related felsic magmas sample at least Mesoproterozoic, Grenvillerelated crust like that of the Hillabee metadacites. This suggests that these arcs sampled a
Laurentian crustal substrate. Only one arc, Exploits, does not have an epsilon Nd value that
involves a large portion of Grenville crust. The GFG also shows similar Nd-isotopic characters
18
like that of Exploits arc. While this is not conclusive in excluding the PCF from being
genetically related to Laurentian crust, it strongly suggests little to no input of Grenville or older
materials incorporated into the magma.
While trying to characterize the trace element content of the arc lavas it might be expected that
the arc lavas represent a mix of three main components, depleted mantle wedge, subducted
sediments and (altered) mafic oceanic crust in case of intra oceanic arc and continental crust in
case of continental margin arcs. It has been recognized for sometime (McCulloch and Gamble,
1991) that a characteristics feature of island arc basalt is the depletion of high field strength
(HFS) elements such as Nb, Ta and to a lesser extent Zr, Ti, Yb, Sc and Ni relative to large ion
lithophile elements (LIL) such as Rb, Ba, Pb, Sr, U and Th. In the PCF amphibolites and
greenstones there is modest depletion of Nb, and Zr, Ti in some samples. There is enrichment in
Ba, Pb in the greenstone which are considered as signatures from dehydrating slab. Pb
enrichment can partly be attributed to the sulfide mineralization in the greenstone. Sr enrichment
in both the amphibolites and greenstones might be due to epidotization. The REE pattern of both
the amphibolites and the greenstones are flat like that of MORB. Considering the flat REE
pattern and modest enrichment of LIL elements with slight Nb a volcanic back-arc source is
preferred. There is good geochemical evidence for transport of LIL enriched volcanic arc melt
into BAB (Saunders and Tarney, 1984; Nakamura et al., 1989).
1.9.2 Tectonic Relationships
On the basis of the new data, and taking into consideration the timing and nature of the
interaction of oceanic and continental tectonic elements, a revised Cambrian –Ordovician
tectonic history of the Laurentian margin is proposed in the southern Appalachians of Georgia
and Alabama. There are still insufficient data to constrain many of the fundamental features of
Laurentian margin tectonic development in the southern Appalachians and there are many
aspects that remain equivocal.
The sequence of events in the beginning of the first Wilson cycle can be summarized as follows:
Grenville orogeny at ca 1 Ga resulting in amalgamation of the Rodinian supercontinent and
extensive intrusion of syn tectonic granites (Karlstrom et al., 2001). Neoproterozoic rifting begun
240 Ma later in the eastern margin of Laurentia with the opening of Iapetus and culminating at
ca. 560 Ma (Cawood et al., 2001). Breakup of Rodinia was multistage (e.g., Colpron et al., 2002
; Harlan et al., 2003) spanning over ~230 m.y. and involving different parts of Laurentia.
Magmatic events that record northward-migrating rift-pulses in the Appalachian Blue Ridge are
the 760–745 Ma Crossnore event in North Carolina (Su et al., 1994 ; Aleinikoff et al., 1995 ).
Other syn rift intrusions during Iapetus opening (e.g., Badger and Sinha, 1988 ; Aleinikoff et al.,
1995) are Robertson River (735– 725 Ma), Battle Mountain events of Virginia Blue Ridge (705–
680 Ma) (Tollo and Hutson, 1996 ; Tollo and Aleinikoff, 1996 ; Tollo et al., 2004 ), Catoctin
eruptive event (572–564 Ma) etc.
The significant thickness of the metasedimentary package of the Ashland-Wedowee belt and
equivalents of the southern Appalachians has caused previous workers to suggest a marginal
basin/slope-rise facies origin for the rocks associated with the Proterozoic rifting of Rodinia,
adjacent to the southeastern margin of Laurentia (Drummond et al., 1994 and 1997). Cawood
and Nemchin (2001) identified 4 distinct zircon populations in the upper Neoproterozoic to
19
Ordovician Laurentian margin sedimentary sequence in the Newfoundland Appalachians that
record a cycle of ocean opening and closing. The four identified groups are as follows (1)
Archean zircons, age range 2850 -2600 Ma. Potential source- Laurentian hinterland. Age
corresponds to zircon crystallization during major magmatic and tectonothermal events in the
Superior craton. (2) Paleoproterozoic zircons, age range 1950 - 1750 Ma. Age corresponds to
craton margin orogenic belts (e.g., Ungava, New Quebec, and Torngat). (3) Mesoproterozoic to
early Neoproterozoic zircons, age range 1450 - 950 Ma. Age corresponds to Grenville orogen
lithologies. (4) Neoproterozoic zircons, age range, 760 - 570 Ma. Age corresponds to numerous
syn-rift igneous intrusions along Laurentian margin, now preserved within the Appalachian
orogen.
The stratigraphic thickness of the Ashland-Wedowee Supergroup in the southern Appalachians
of Alabama and Georgia reaches >10 Km at places. The lithologic association of the AshlandWedowee Supergroup, predominantly deep water turbidite metasediments and subordinate
intercalated tholeiitic rift/ocean floor basalts points towards a significant sediment source. The
possibilities include 1. the sediments represents slope-rise sequence of a rifted continental
margin that prograded out on the oceanic crust (now preserved as the tholeiites) 2. intra oceanic
complex. The later option is improbable due to the relative proportion of sediments to basalt. The
Ashland-Wedowee Supergroup is in fault contact and tectonically lies immediately above the
Laurentian outer shelf rocks. Presence of Middle Proterozoic zircon xenocrysts in numerous
plutons (product of anatectic melts of EBR units) that intrude the Ashland-Wedowee sediments
indicates that Alabama EBR formed near or at the Laurentian rifted margin. It has been
suggested that the Ashland-Wedowee Supergroup sediments qualifies as the best and possibly
only candidate for the late Precambrian clastic slope/rise outer margin prism along the Alabama
promontory during Rodinia rifting (Tull, 1978; Thomas et al.,1980; Stow et al., 1984;
Drummond et al.,1988; 1994; 1997).
In case of the Ashland-Wedowee group sediments because there are no distinguishing
characteristics like fossil assemledge or basement clasts to ascertain the age and composition of
the basement beneath the Ashland Supergroup, Nd model ages (based on CHUR extraction) were
computed to determine the source characteristics of the Ashland-Wedowee Belt sediments.
Majority of the samples exhibited ages between 950-1150 Ma, with one outlier at 1440 Ma. The
ages are typically of Laurentian Grenville basement (the older age is similar to the GraniteRhyolite province) and suggests that Ashland-Wedowee belt sediments were supplied from
Laurentian margin.
Increasing Nd isotopic data now exist in the literature for crystalline rocks, specially the
granitoids of the Blue Ridge. The Nd isotopic data for the metasediments overlap the field
defined for the Grenville basement data (recalculated at 700 Ma, approximate time of deposition
of the Ashland metasediments) defined by Carrigan et al., (2003) and Hatcher et al., (2004) (Fig.
1.17). The oldest TDM age reported by Bream et al., (2004) is 1.9 Ga from the Dahlonega Gold
belt. Eastern Blue Ridge metagraywacke and quartzites from Tallulah Falls Formation ranged in
age between 1.24 to 1.61 Ga (Bream et al., 2004). The TDM ages of the metasediments obtained
from Ashland Supergroup and Canton Schist agrees well with the previous data.
20
The detrital zircon data is obtained from one sample within the Pumpkinvine Creek formation
that shows a pronounced age cluster between 918 to 1190 Ma which is similar to the Grenville
magmatic event. Few pre-Grenville components like Middle Proterozoic grains (1351-1462 Ma)
are also present. In the absence of detrital zircon data on the Ashland Wedowee samples it is
difficult to interpret the old age of 1892 Ma but middle Proterozoic and late Archean grains have
been identified from eastern Blue Ridge sediments and the Dahlonega Gold belt sediments by
Bream et al., 2004. Majority of the detrital zircons analyzed from Dahlonega Gold belt
metasandstones reflect a North American source except for one sample that contains Gondwana
(?) (1.8-2.0 Ga) zircons, in addition to Grenville and older North American detrital zircons
(Bream, 2003). The Gondwana ages have not been reproduced in the other analyzed samples
(Bream, 2003; Hatcher et al., 2006). The zircons in the one sample with Gondwana affinity could
have been derived from the Laurentian Penokean orogen that occurred during the same time.
During Rhodinia supercontinent rifting, numerous crustal fragments and an array of complex
depositional basins between Laurentia and the South American craton was created. Within the
southern Appalachian crystalline core there is exists strong evidence for both Laurentian and
exotic sediment sources. The predominating detrital zircon population is aged between 1.0–1.25
Ga from the metasediments of southern Appalachian that were inboard of the Carolina terrane.
This evidence suggests that these basins were placed near a Grenvillian source. This
interpretation is consistent with the nonconformities that exist between many basement and cover
contacts in the Blue Ridge.
Orthoamphibolites commonly interlayered with graphite-quartz schist and quartzite and garnetquartzite/schist are common in the upper part of the Ashland Group. They range in thickness
from centimeters to several 100 m and make up about 7% of these equivalent groups (Tull et al.,
2006). A volcanic or plutonic protolith for the amphibolite is suggested from field relationships
(interlayered, locally gradational, concordant contacts with surrounding metasediments) and
textural evidence (orthopyroxene relics, surrounded by coronas of hornblende). Geologic setting
combined with geochemical, textural, and mineralogic criteria led geologists to interpret these
rocks to be tholeiitic ocean-floor basalts intercalated with deep water metasediments. (Tull,
1978; Thomas et al., 1980; Stow et al., 1984; Drummond et al., 1988). Small lenses and pods of
ultramafic rocks are locally present in the amphibolite (Reynolds, 1973; Mies and Dean, 1994).
Lithologic contacts between this thick pile of amphibolites with intercalated metasediments
indicates that they are not faulted, but are interlayered with metasedimentary and other possible
metavolcanic lithologies. There is complete lack of any primary facing data in these units, but it
is highly unlikely that this >10 km thick sequence is regionally overturned. It is suggested that
the amphibolites originated as ocean floor tholeiites and are related to the initial rifting of the
supercontinent Rhodinia.
Celar Sengor and Burke, 1978 described the process of evolution of continental rifts can occur in
two ways: active rifting or passive rifting. 1. Active rifting – it is initiated by upwelling of
magma in the asthenosphere thus often called mantle-generated rifting. In case of active rifting
the zone of uplift and the zone of thinning extend for many kilometers far beyond the rift margin.
2. Passive rifting – it is caused by extensional forces in the lithosphere thus referred to as
lithosphere-generated rifting. The zone of uplift in passive rifting and the zone of thinning are on
the surface of the earth and confined to the rift margin. Fig. 1.23 summarizes the common
21
concepts conceived for passive extension models. Plate tectonics is the primary driving force for
lithospheric extension in the passive model. The buoyant sub-lithospheric mantle passively
upwells beneath the thinned area (McKenzie, 1978). Subsidence followed by sedimentation
occurs continuously, both during syn-rift (initial subsidence) and post-rift (long-term subsidence)
stages. Thus mantle melting in this case is a late consequence of the dynamic stretching of
lithosphere (White and McKenzie, 1989).
The amagmatic rifted margins (ARM) typically have a weak crust dominated by faults and
abundant serpentinite. Different rift margins are characterized by difference in magma budgets
(extremely low magmatism for the ARM). Thus the lithosphere beneath also vary from ‘depleted
mantle beneath Volcanic Rifted Margins and undepleted mantle beneath ARMs’ (Stern, 2004).
Though Cenozoic examples of passive rifted margins are rare but there are several well
documented passive margins initiated rifting during the Mesozoic time. Examples include the
west Iberia-Newfoundland conjugate margins and Alpine Tethyan margins. In the former case
the final rifting phase and sea floor spreading was during Early Cretaceous (~133Ma)
(Whitmarsh et al., 2001) and the later experienced the final phase of rifting during Triassic/Early
Jurassic presaging the opening of the Liguria-Piemonte ocean where sea floor spreading dates
back to 160- 165Ma ago (Whitmarsh et al., 2001). Both the above mentioned cases are
characterized by magma-poor margins with margin-parallel deep-water zones that indicates
successive stages of margin evolution. The continental crust along the margin is thinned and
dissected by numerous low angle detachment faults. Sub-continental mantle exhumation follows
faulting and typically the mafic melts volumetrically increases oceanward and merges with the
oceanic crust (Dean et al., 2000; Boillot et al., 1989).The current lithologic association and
stratigraphic thickness of the Ashland-Wedowee belt syn-post rift sediments and the presence of
a very low percent of amphibolites might be related to the initiation of passive margin rifting of
the Rhodinia supercontinent during the Neo Proterozoic time at least in the southern part of the
Appalachians.
The occurrence of granitic intrusion is widespread within the Ashland Wedowee belt that were
emplaced between ~490-430 Ma (Russell, 1978; Russell et al., 1987; Higgins et al., 1997).
Drummond (1997) suggests that the oldest intrusive Elkahatchee Quartz Diorite (EQD) formed
by subduction and melting of a westward dipping oceanic slab beneath Laurentia during the
earliest Ordovician. The younger granitoids (like Rockford type) are thought to accompany
discrete dynamothermal metamorphic events (Russell, 1978). Zircons from the Ordovician-early
Silurian Austell (Higgins et al. 1997) and Mulberry Rock gneisses include a fraction of inherited
Grenville aged (~1.1 Ga) zircons and zircon cores suggesting that these formed while
incorporating or intruding into Laurentian-affinity, Grenvillian crust. Cambrian to Late
Ordovician plutonic bodies within the EBR likely formed the core of a continental margin arc
(Drummond et al., 1997). Tull et al., (2006) suggested that as the Proterozoic- EBR strata served
as the arc basement of subduction-generated magmas associated with a continental arc and the
present erosional level would place this sequence several kilometers beneath the former arc
volcanic carapace. If a ~460 Ma Hillabee and PCF were cogenetic with the plutonic complexes,
then it could have constituted part of that overlying volcanic carapace in an arc to back-arc
position. If the main part of the volcanic arc overlay the current EBR it would have been stripped
by erosion and be absent at the current erosion level. A cogenetic association of the Hillabee and
PCF with the EBR plutonic arc core suggests that B-type subduction may have spanned a period
22
of ~35 m.y., based upon the ages of the plutons. A similar span may also be reflected in the
GFG and Hillabee dacite zircon data with Cambrian core at ~500 Ma., i.e. 30-40 Ma older than
the interpreted GFG and Hillabee dacite crystallization age. Thus, in modeling subduction
events related to a cogenetic EBR and PCF-Hillabee, the time span of subduction is assumed to
be ~30-40 Ma.
It has been suggested by previous workers (e.g. McConnell, 1980; McConnell and Abrams,
1984; Higgins et al., 1988; McClellan et al., 2005; Holm and Das, 2005) that the Hillabee
Greenstone of Alabama and the PCF are equivalent units. The Hillabee and the PCF have very
similar geochemical characteristics and occupy a similar structural position adjacent to the
Talladega Belt-western Blue Ridge. Geochemical, isotopic and geochronological evidence
suggests that the PCF and Hillabee are both Ordovician volcanic arc/back arc that marked the
beginning of Taconic orogeny in the southern Appalachians. Both the GFG and the Hillabee
dacite lack any inherited Grenville zircon but the ε Nd values for the dacites are more negative
that the GFG indicating some interaction with older crust. The negative ε Hf values (Tull,
personal communication) of the Hillabee dacite zircons attests the interaction of the Hillabee
with an older crust. If the PCF and Hillabee are parts of the same arc-back arc system then the
subduction zone might have stepped across a transform fault. In that case PCF were formed due
to subduction against an oceanic crust where as the Hillabee formed on the continental crust. The
background sedimentation that accompanied the arc building was receiving detritus from
Laurentia as evidenced by the Nd model age of the Canton schist and detrital zircon of Grenville
affinity in the meta-sandstone of PCF. However, there is no evidence of deformation along the
southeastern Laurentian margin until the Devonian-Mississippian. Thomas (2004) suggests that
because there is no deformation in the palinspastic location of the Talladega belt during the
Taconic or Acadian orogenies and that the accreted terranes remained at or southeast of the rifted
margin of Laurentia until accretion during the Alleghenian orogenic event.
The Hillabee is in thrust contact with the Talladega Group and has the same metamorphic grade
and concordant metamorphic fabric. Thus the time of thrusting is thus bracketed by the age of
the uppermost member of the Talladega Group and metamorphic age. The table (Table I) below
summarizes the available age data and their possible interpretations regarding the metamorphic
age of the EBR and thrusting age of HG on Talladega belt.
The (327-333) Ma 40Ar/39Ar biotite, hornblende ages of the amphibolite facies EBR rocks are
interpreted to be metamorphic age of Talladega belt caused by emplacement of EBR over
Talladega belt (Kish, 1990, Steltenpohl et al., 2005). This interpretation seems unlikely because
if EBR is emplaced directly on cold Talladega belt sediments the hot basal part of the EBR
would show evidence of extensive retrograde metamorphism which is not observed in the field
(Tull et al., 2006).
23
Table I : Available age data and their possible interpretations regarding the metamorphic age of
the EBR and thrusting age of HG on Talladega belt.
Material dated
Age
Dating
References
Possible age interpretation
technique
Upper
Erin 245-375
Fossil
Tull et al., Deposition age of Erin shale
Shale
Ma
assemblage
1989;
and
the
youngest
Gastaldo,
stratighaphic age of the
1995
Talladega belt.
Talladega belt ~370-400
K/Ar, Rb/Sr Wampler et The oldest age (older than
slate ages, some Ma
and 40Ar/39Ar al.,
1970; the age of the Erin shale)
from Erin shale.
whole- rock Tull,
1982; possibly is from detrital Kslate ages
Kish, 1990; bearing minerals in the
Durham, 1993 metapelites (Kish, 1990).
40
39
Talladega Belt 321-334Ma
Ar/ Ar
McClellan et Cooling age and not a peak
metasedimentary
analysis
of al., 2005
metamorphic age.
and metaigneous
white micas.
rock
Hornblende and 333.8 ±1.7 40Ar/39Ar ages Kish, 1990, Metamorphic
age
of
Biotite
from Ma
Steltenpohl et Talladega belt caused by
EBR
al., 2005
emplacement of EBR over
327.4 ±1.6
amphibolite
Talladega belt.
Ma
facies rocks.
Syn to post ~ 366-370 U-Pb zircon ages Russell, 1978; Peak metamorphic age of
and
Rb/Sr Russell et al., EBR.
kinematic
Ma
whole rock ages.
1987;
igneous
Steltenpohl et
intrusions
in
al., 2005
EBR.
Similar metamorphic ages obtained from EBR and Talladega Belt suggest that metamorphism
and accretion of the Hillabee Greenstone, and the PCF, were synchronous with and completed
during the Alleghenian orogeny.
Palinspastic restoration of the paleotectonic position of the Alabama/Georgia eastern Blue Ridge
places it eastward to the present location of the Pine Mountain internal basement massif. This has
been done by strike perpendicular restoration of Talladega belt carbonate shelf strata and lower
Paleozoic foreland (Thomas, 2004). A Laurentian affinity of the Ashland-Wedowee belt
sediments makes it a suitable candidate for the clastic continental slope/rise outer margin prism
that was forming over the Neoproterozoic – Cambrian margin of southeastern Laurentia. From
the available stratigraphic and structural data and in the light of the new geochemical and
geochronological data the following constrains are placed while trying to predict a model for the
origin and evolution of the eastern Blue Ridge, southern Appalachian: A) the Ashland –
Wedowee sediments is interpreted as slope rise deposit adjacent to Laurentia, B) the HG and
PCF formation are part of the same Ordovician suprasubduction arc-back arc system probably
involving Mesoproterozoic continental crust C) HG currently lies directly above Laurentian
outer shelf strata whereas PCF formation is a part of the eastern Blue Ridge sequence D)
24
Mulberry rock gneiss is one of the numerous Ordovician-Silurian pluton and is not a Grenville
basement in the Talladega belt. The geochemical signature that the HG and PCF formed as part
of a continental arc system, their present structural position, and age similarity of the eastern
Blue Ridge plutonism argue strongly that the HG-PVC is native to Laurentia and was not part of
an exotic island arc. If the PCF is equivalent to the Greenstone, then accretion must take place
after the Devonian/early Mississippian and prior to metamorphism constrained by 40Ar/39Ar ages
of 334-320 Ma (Barineau and Tull, 2001; McClellan et al. 2005). The PCF and Hillabee, exist
only as fragments, specific to the back-arc region, unlike the Taconic arcs of the New England
Appalachians that commonly include typical arc related features such as voluminous arc-related
sediments, accretionary wedge, and evidence of accretion in the foreland. Figure 1.24
summarizes the tectonic model proposed for the southern Appalachians during the Ordovician
Taconic Orogeny in this study. If the PCF and Hillabee are the remnants of the Ordovician
volcanic arc, the tectonic setting, timing and style of the Taconic orogeny in southern
Appalachians and formation of the arc-related terranes took place in an environment very
different from the Ordovician arcs of the northern and central Appalachians.The arc in the
southern Appalachian formed during Ordovician but remained outboard of the Laurentian
margin until terminal closure of the Iapetan Ocean due to continent-continent collision between
Laurentia and Africa.
25
Table 1.1. Major Oxide concentration in weight percent.
Sample
Unit
SiO2
Al2O3 Fe2O3T MgO
M68
M201B
M243
M268
M271
M272
SM1100
M67
D1
D2
BG995
BG999
Y 23 C
Y 24
Y 131
NT 911
BH 166
SC 318G
TA 7
US 41A
Y 126
Y 172
TA 16
TA 17
SC 254
SC 318A
AC 78
BH 57
G
G
G
G
G
G
G
Metadacite
Metadacite
Metadacite
Metadacite
Metadacite
GFG
GFG
GFG
GFG
GFG
GFG
GFG
GFG
PCF
PCF
PCF
PCF
PCF
PCF
PCF
PCF
47.46
47.07
48.04
49.24
51.04
51.21
48.05
67.71
66.45
66.32
65.70
65.90
75.28
73.64
76.75
74.13
78.11
76.29
73.88
73.79
50.29
49.12
47.79
50.47
43.85
50.26
48.51
48.34
16.44
15.57
15.40
14.17
13.88
14.76
13.90
15.44
16.30
16.34
16.90
16.74
13.67
14.31
14.56
14.83
13.37
12.61
14.19
12.83
15.12
15.63
17.19
14.93
17.67
14.37
14.47
16.05
CaO
10.70 8.53 13.58
12.54 10.07 10.74
12.93 7.55 11.26
11.73 8.51 12.42
13.15 6.44 10.93
9.21 8.28 12.31
10.75 11.28 14.21
4.03 1.89 4.26
3.95 1.81 3.63
4.17 1.97 3.56
3.47 1.87 3.62
3.89 1.62 3.89
1.98 0.40 1.62
3.95 0.71 1.39
2.13 0.29 0.13
2.66 0.47 1.34
2.05 1.28 0.86
2.90 1.22 0.74
2.39 0.70 2.49
3.85 1.43 1.43
11.37 7.66 11.46
13.97 4.28 10.00
9.60 8.56 13.63
11.60 6.96 11.20
14.01 4.98 16.11
13.18 7.76 9.10
12.47 7.69 12.78
11.24 9.26 10.70
Na2O K2O TiO2 P2O5
MnO % LOI
1.77
2.02
2.86
2.36
2.31
3.11
1.16
3.29
3.97
4.12
3.98
3.76
6.63
4.18
3.80
3.93
3.64
6.00
4.59
5.51
2.57
3.82
1.93
3.06
1.28
3.79
2.32
2.86
0.195
0.166
0.185
0.164
0.204
0.154
0.196
0.061
0.070
0.060
0.060
0.070
0.020
0.071
0.031
0.030
0.010
0.030
0.051
0.071
0.162
0.184
0.145
0.163
0.223
0.173
0.172
0.154
Wt % determined by ACME laboratory by ICP- emission spectrometry.
Wt % corrected for LOI (Loss of ignition)
Total iron reported as Fe2O3T
G - Hillabee Greenstone
Metadacite – Hillabee dacite
GFG - Galts Ferry Gneiss
PCF - Pumpkinvine Creek Amphibolites
26
0.05
0.22
0.05
0.04
0.15
0.07
0.04
2.77
2.88
2.93
3.37
3.12
0.26
1.42
1.95
2.30
0.46
0.04
1.17
0.35
0.14
0.15
0.07
0.06
0.07
0.04
0.06
0.05
0.98
1.22
1.35
1.18
1.54
0.68
0.37
0.36
0.45
0.54
0.39
0.41
0.14
0.27
0.21
0.22
0.19
0.23
0.20
0.27
1.07
2.39
0.94
1.34
1.47
1.18
1.18
1.06
0.082
0.114
0.123
0.092
0.153
0.041
0.021
0.091
0.080
0.110
0.100
0.080
0.010
0.061
0.010
0.051
0.031
0.020
0.051
0.031
0.101
0.347
0.062
0.142
0.203
0.081
0.121
0.113
2.1
1.8
1.8
1.6
1
1.6
1
1.9
1.0
1.3
0.9
1.2
1
1.3
1.8
1.3
2.4
1.6
1.2
1.8
1
2
2.4
1.6
1.5
1.6
1.2
2.9
Table 1.2. Trace Element and REE concentration (in ppm) of the Pumpkinvine Creek
Amphibolites (PCF) and the Galts Ferry Gneiss (GFG).
Y 126 Y 172 TA 16 TA 17
Unit
Li
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Yb
Lu
Hf
Ta
Pb
Th
U
Sc
Ti
V
Cr
Co
Ni
Cu
Zn
Ga
SC
SC AC 78 BH 57
254 318A
PCF PCF PCF PCF PCF PCF PCF PCF
_
_
_
_
_
_
_
_
2.60 1.80 2.40 0.50 0.80 2.10 2.70 0.90
165
327
200
198
646
106
223
229
31.6 50.8 19.8 30.5 43.9 27.4 29.9 25.2
66.4 188.4 51.2 81.6 100.3 67.8 69.0 57.2
3.3
9.3
2.5
4.4
4.1
2.1
2.8
2.6
0.10 0.08 0.20 0.05 0.20 0.13 0.20 0.18
14.9 25.7 59.5 13.7
8.6
12.3 13.9 40.2
4.9
12.6
3.6
5.6
7.3
3.9
3.8
3.1
12.5 29.9
8.6
12.8 14.6
9.1
9.8
8.3
1.89 4.07 1.28 1.99 2.40 1.47 1.53 1.26
10.20 21.70 6.70 10.90 13.40 8.30 8.60 7.40
3.50 6.10 2.40 3.50 4.20 3.00 2.80 2.60
1.05 2.31 0.83 1.24 1.62 0.85 1.09 0.85
4.22 7.35 2.99 4.38 5.46 3.82 4.16 3.44
0.78 1.26 0.50 0.77 1.09 0.70 0.72 0.66
4.99 8.22 3.19 5.04 6.73 4.37 4.87 4.19
1.05 1.64 0.64 1.07 1.43 0.94 1.02 0.86
3.02 4.84 1.94 2.95 4.36 2.71 3.02 2.39
3.08 4.74 1.90 2.97 4.56 2.63 2.73 2.54
0.46 0.76 0.28 0.43 0.71 0.44 0.45 0.39
2.00 5.60 1.70 2.30 3.00 2.00 2.30 1.90
0.30 0.60 0.30 0.30 0.50 0.20 0.20 0.20
0.20 0.80 0.10 0.30 0.50 0.20 0.40 0.20
0.50 1.00 0.30 0.40 0.60 0.30 0.10 0.30
0.20 0.40 0.10 0.10 0.60 0.10 0.10 0.10
41
35
36
41
42
44
47
49
6419 14315 5648 8042 8825 7067 7099 6359
258.0 371.0 232.0 315.0 479.0 314.0 326.0 295.0
_
_
_
_
_
_
_
_
43.6 32.0 39.0 48.0 49.7 55.6 57.5 74.1
74
22
94
35
51
43
79
98
11.6 109.1 31.9 59.9 28.6 73.6 105.9 150.3
9
29
18
20
9
14
10
27
17.3 25.8 16.3 19.3 29.1 16.5 20.0 17.5
27
Y23C
Y24
Y 131
NT
BH
SC
TA7 US41
911
166 318G
GFG GFG GFG GFG GFG GFG GFG GFG
5.36 4.72 1.41 4.64 7.96 1.00 2.52 5.33
13.90 39.68 60.73 75.75 12.96 0.40 33.25 10.52
58
89
64
129
99
25
194
48
25.0 14.4
8.4
9.8
11.2
9.1
8.5
34.9
94.6 44.6 75.4 45.6 26.7 33.8 57.0 44.9
7.6
6.2
5.9
5.3
1.7
3.5
3.2
5.0
0.38 0.27 0.40 0.72 0.13 0.04 0.26 1.05
53.0 336.1 461.0 627.6 115.8 15.5 399.5 64.9
32.0
8.4
4.7
9.7
5.6
2.8
20.0 10.7
64.6 21.6 35.9 14.1 14.7
6.5
38.6 24.6
7.10 2.44 1.06 1.88 1.44 0.94 3.07 3.04
24.72 9.53 3.69 6.72 5.70 4.31 9.42 12.92
4.62 2.27 0.82 1.35 1.28 1.15 1.48 3.42
0.51 0.58 0.31 0.55 0.28 0.37 0.58 0.71
6.18 2.35 2.47 1.60 1.46 1.03 3.39 3.23
0.72 0.40 0.18 0.23 0.23 0.21 0.24 0.69
4.07 2.78 1.36 1.44 1.60 1.53 1.21 5.27
0.84 0.59 0.31 0.31 0.35 0.32 0.25 1.15
2.59 1.80 1.00 1.00 1.14 0.99 0.80 3.65
2.48 1.74 1.02 1.03 1.11 1.01 0.84 3.47
0.36 0.27 0.17 0.17 0.18 0.17 0.14 0.53
2.66 1.43 2.21 1.37 0.83 0.83 1.57 1.37
1.09 0.70 0.65 0.56 0.32 0.37 0.88 1.03
4.28 20.59 19.14 20.13 8.04 0.92 6.79 2.08
15.32 2.85 8.44 4.17 2.47 0.89 9.42 1.69
6.47 1.95 2.45 3.44 0.90 0.21 4.97 0.58
4.76 4.71 4.39 3.43 6.91 15.36 4.96 10.43
862 1992 1302 1441 530 1348 1219 1618
17.2 23.4 12.1 29.2 21.1
3.6
41.9 19.9
0.71 1.80 1.67 2.51 2.99 1.05 3.61 2.98
10.8 20.9 12.2 11.2 16.8 12.1 53.4 51.0
0.1
2.7
0.7
2.4
2.5
3.5
2.2
0.7
23.3 10.4
5.2
1.4 156.9 2.5
0.9
1.4
11.6 50.3 24.9 21.2 17.4 40.8 14.5 45.0
1.7
10.7 14.1 19.6
3.6
0.5
12.1
2.0
Table 1.3. Trace Element and REE concentration (in ppm) of the Hillabee Greenstone (G)
& Hillabee dacite (dacite).
Unit
Li
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Yb
Lu
Hf
Ta
Pb
Th
U
Sc
Ti
V
Cr
Co
Ni
Cu
Zn
Ga
M68 M201b M243
G
G
G
6.89 12.44 9.17
0.89
3.52
0.40
106
180
153
19
23
31
64
53
62
3.28
3.33
4.45
0.12
0.05
0.03
33
26
235
3.05
4.61
8.52
7.91 10.32 14.81
1.23
1.72
3.15
6.46
8.51 14.51
2.06
2.63
4.14
0.75
0.98
1.42
1.69
2.16
3.09
0.44
0.53
0.80
3.35
4.04
5.88
0.71
0.83
1.19
2.05
2.44
3.41
1.77
2.08
2.85
0.26
0.29
0.37
0.65
0.24
0.73
0.27
0.23
0.33
1.04
0.65
1.20
0.25
0.24
0.58
0.07
0.09
0.17
39.7
35.3
46.2
5734 6864 7831
240
235
311
565
781
506
47
59
57
116
299
100
60
66
126
73
79
82
14
16
16
M268
G
4.71
0.52
134
24
73
3.52
0.06
154
4.16
10.44
1.57
8.03
2.56
0.97
2.09
0.53
4.06
0.86
2.50
2.06
0.29
0.85
0.29
0.83
0.41
0.12
41.1
6763
273
513
54
151.
74
77
14
M271
G
8.84
6.02
140
33
52
6.82
0.32
24
8.97
21.26
3.02
14.05
3.97
1.20
3.55
0.79
5.59
1.16
3.36
2.82
0.41
1.58
0.47
4.14
1.64
0.38
41.6
8985
319
103
55
55
81
105
17
M272 Sm1100
G
G
6.08
4.53
0.75
0.27
106
64
15
8
58
56
0.84
0.35
0.03
0.05
23
36
1.21
0.82
3.64
2
0.61
0.31
3.48
1.71
1.29
0.66
0.54
0.35
1.01
0.55
0.31
0.17
2.54
1.42
0.55
0.31
1.68
0.92
1.50
0.82
0.23
0.13
0.44
0.26
0.11
0.09
1.02
0.38
0.13
0.09
0.04
0.03
38.6
39.2
3950
2117
240
162
481
243
41
67
119
111
58
145
63
52
14
10
28
M67
D1
D2
BG995
BG999
dacite
dacite
dacite
dacite
dacite
12.65
123
226
19
71
10.55
2.10
337
18.81
42.30
5.30
19.56
4.41
1.02
4.69
0.64
3.62
0.69
2
1.91
0.28
2.47
1.14
13.74
12.22
4.63
10.8
2085
63.2
49.4
20.2
16.7
1.9
41.6
18
15.26
158
181
7
455
5.72
3.49
530
19.43
39.76
4.18
13.43
2.12
0.53
1.02
0.17
1.37
0.28
0.84
0.73
0.14
12.81
0.60
6.87
9.04
1.94
5.0
806
26.0
40.4
4.5
16.7
1.5
16.2
3.6
6.78
127
181
14
537
8.51
2.94
478
25.98
53.64
5.76
20.04
3.53
0.59
1.48
0.28
2.51
0.50
1.50
1.25
0.23
16.45
0.73
10.04
11.78
2.67
7.2
1138
31.7
129.8
5.9
28.4
56.5
19.4
3.4
11.55
125
195
11
504
6.95
2.34
429
22.06
47.04
4.84
16.64
2.96
0.54
1.27
0.24
2.10
0.42
1.26
1.06
0.19
15.32
0.56
10.20
12.12
2.40
7.2
1002
32.5
122.1
5.7
25.8
45.4
20.7
3.3
8.95
163
184
11
650
6.45
3.13
441
24.19
48.56
5.13
17.21
3.01
0.53
1.28
0.24
2.16
0.44
1.32
1.12
0.20
20.15
0.64
7.53
13.34
3.28
6.9
958
29.3
104.7
5.7
20.9
149.4
20.9
3.4
Table 1.4. Sr and Nd Isotopic Data.
Sample
Units
M68
G
M201b
G
Sm
Nd
Rb
Sr
(ppm) (ppm) (ppm) (ppm)
2.06
6.46
0.89 106.12
2.63
8.51
3.52
180.43
147
Sm/144Nd
143
Nd/144 Nd
87
Rb/86Sr
87
Sr/86Sr
(87Sr/86Sr)I εNd present εNd 460
0.1970
0.512939 (10)
0.0245
0.705891 (09)
0.705732
5.87
5.86
0.1911
0.512932 (08)
0.0571
0.705639 (09)
0.705270
5.74
6.07
M243
G
4.14
14.51
0.40
153.05
0.1760
0.512717 (15)
077
0.705932 (08)
0.705882
1.54
2.76
M268
G
2.56
8.03
0.52
133.95
0.1971
0.512923 (12)
0.0113
0.705118 (12)
0.705045
5.56
5.54
M271
G
3.97
14.05
6.02
139.61
0.1745
0.512674 (10)
0.1263
0.706134 (12)
0.705317
0.70
2.01
M272
G
1.29
3.48
0.75
105.74
0.2291
0.513086 (09)
0.0207
0.706485 (11)
0.706351
8.74
6.84
Sm1100
G
0.66
1.71
0.27
64.11
0.2394
0.51302 (09)
0.0125
0.703986 (09)
0.703905
7.45
4.95
M 67
HD
4.41
19.56
123.1
226.1
0.1392
0.512192 (10)
1.5937
0.718392 (08)
0.708080
-8.70
-5.55
D1
HD
2.12
13.43
158.4
180.7
0.0972
0.512103 (10)
2.5671
0.723548 (11)
0.706939
-10.44
-4.59
D2
HD
3.53
20.04
127.2
180.8
0.1088
0.512136 (12)
2.0612
0.721219 (18)
0.707883
-9.79
-4.63
BG995
HD
2.96
16.64
125.5
195.2
0.1100
0.512146 (14)
1.8829
0.720451 (07)
0.708269
-9.60
-4.51
BG999
HD
3.01
17.21
163.1
184.3
0.1080
0.512169 (13)
2.5910
0.724355 (09)
0.707591
-9.15
-3.94
Y 126
PCF
3.50
10.20
2.60
164.6
0.2118
0.51296 (11)
0.0463
0.70441 (14)
0.704106
6.28
5.33
6.64
Y 172
PCF
6.10
21.70
1.80
326.7
0.1735
0.512908 (10)
0.0161
0.70474 (12)
0.704631
5.27
TA 16
PCF
2.40
6.70
2.40
200.1
0.2211
0.512937 (16)
0.0351
0.70459 (10)
0.704360
5.83
4.40
TA 17
PCF
3.50
10.90
0.50
198.3
0.1982
0.512938 (24)
074
0.70413 (10)
0.704080
5.85
5.77
6.38
SC 254
PCF
4.20
13.40
0.80
646.6
0.1935
0.512955 (14)
036
0.70693 (09)
0.706906
6.18
SC 318A
PCF
3
8.30
2.10
106
0.2231
0.512887 (12)
0.0580
0.70579 (10)
0.705419
4.86
3.31
AC 78
PCF
2.80
8.60
2.70
223.3
0.2009
0.513045 (10)
0.0354
0.70666 (12)
0.706430
7.94
7.70
BH 57
PCF
2.60
7.40
0.90
229.1
0.2169
0.513029 (17)
0.0115
0.70482 (13)
0.704746
7.63
6.45
Y 23 C
GFG
5
26.40
13.90
57.68
0.1169
0.512233 (09)
0.7055
0.712751 (08)
0.708186
-7.90
-3.21
Y 24
GFG
3
12.10
39.68
88.70
0.1530
0.512510 (09)
1.3098
0.714149 (08)
0.705674
-2.50
0.07
Y 131
GFG
1.10
5.10
60.73
63.46
0.1331
0.512379 (10)
2.8025
0.722402 (07)
0.704270
-5.05
-1.32
NT 911
GFG
2.40
12.20
75.75 128.82
0.1214
0.512382 (11)
1.7220
0.71674 (09)
0.705598
-4.99
-0.57
BH 166
GFG
1.40
6
12.96
99.32
0.1440
0.512514 (10)
0.3821
0.710751 (07)
0.708279
-2.42
0.68
SC 318G
GFG
1.60
4.80
0.40
24.77
0.2057
0.512750 (12)
0.0472
0.706339 (10)
0.706034
2.18
1.66
TA 7
GFG
1.30
10.10
33.25 193.69
0.0794
0.512257 (16)
0.5027
0.710996 (10)
0.707743
-7.43
-0.54
US 41A
GFG
3.80
14.20
10.52
0.1652
0.512781 (09)
0.6465
0.709467 (09)
0.705284
2.79
47.67
4.65
εNd 438
MR 1
MRG
5.5
27.1
0.1253
0.512292(10)
-6.75
-2.83
MR 3A
MRG
5.7
20.8
0.1691
0.512449(10)
-3.69
-2.18
MR 2a
MRG
3.1
16.2
0.1181
0.512321(09)
-6.18
-1.87
MR 2b
MRG
3.7
18.7
0.1221
0.512285(12)
-6.89
-2.79
MR 3b
MRG
6.5
25.5
0.1573
0.512432(16)
-4.02
-1.86
MR 4a
MRG
6.5
23.0
0.1744
0.512462(12)
-3.43
-2.21
MR 4b
MRG
5.8
23.1
0.1550
0.512435(09)
-3.96
-1.67
MR 4c
MRG
6.2
22.6
0.1693
0.512461(13)
-3.45
-1.95
Type of units:
G - Hillabee Greenstone HD – Hillabee dacite PCF - Pumpkinvine Creek Amphibolites GFG - Galts Ferry Gneiss
MRG – Mulberry Rock Gneiss
2σ errors on 143Nd/144 Nd and 87Sr/86Sr ratios reported as last two significant digits. 143Nd/144 Nd measured ratio normalized
to 146Nd/144 Nd = 0.7219. 143Nd/144 Nd present day = 0.512638 and 147Sm/144 Nd present day = 0.1967
Initial ε values were calculated from the present day 143Nd/144NdCHUR = 0.512638 and 147Sm/144NdCHUR = 0.1967.
Initial 87Sr/86Sr values were calculated for present day bulk earth value of 87Sr/86Sr = 0.7047, and 87Rb/86Sr = 0.0847.
29
Table 1.5. Rb-Sr isotopic data for whole rock samples.
Sample
MR 1
MR 3A
MR 2a
MR 2b
MR 3b
MR 4a
MR 4b
MR 4c
Y 23 C
Y 24
Y 131
NT 911
BH 166
SC 318G
TA 7
US 41A
Unit Rb (ppm) Sr (ppm)
MRG
201
106
MRG
315
24.2
MRG
219
53.2
MRG
262
19.5
MRG
312
28.6
MRG
309
18.5
MRG
277
25.5
MRG
315
17
Rb/Sr
1.90
13.03
4.11
13.41
10.91
16.68
10.87
18.55
87
GFG
GFG
GFG
GFG
GFG
GFG
GFG
GFG
0.227
0.437
0.864
0.447
0.127
0.019
0.163
0.200
0.71275
0.71415
0.72240
0.71674
0.71075
0.70634
0.71100
0.70947
14.1
47.5
59.8
86.7
13.1
0.5
32.8
10.6
62.2
108
69.2
194
103
26.8
201
53
Sr/86Sr
0.74242
0.78318
0.96096
0.95364
0.91850
1.01209
0.91588
1.04346
87
Rb/86Sr
5.13
11.76
38.49
36.40
32.12
44.58
30.33
51.95
0.7136
1.4265
2.7073
0.3500
0.0483
0.4674
0.5991
4.3788
Type of units:
GFG - Galts Ferry Gneiss
MRG – Mulberry Rock Gneiss
87
Sr/86Sr measured ratio normalized to 86Sr/88Sr = 0.1194.
Precision of replicate whole rock determinations is 0.01% for 87Sr/86Sr.
87
Rb/86Sr ratio of MRG determined by isotope dilution and that of GFG measured by ICP-MS.
Rb and Sr concentrations determined by ICP-MS.
Table 1.6. Nd Model Age of metasediments from eastern Blue Ridge.
Sample
Lithology
Sm
(ppm)
3.73
3.62
8.06
2.48
7.10
5.65
5.96
4.34
7.20
8.92
3.89
15.52
13.65
17.01
7.49
7.56
Nd
(ppm)
20.03
19.41
41.48
12.51
43.80
29.39
32.19
21.65
36.72
48.09
18.67
87.85
78.27
101.64
40.03
40.84
147
Sm/144Nd
143
Nd/144 Nd
TCHUR TDM3.4
(Ma) (Ma)
1048 1545
1055 1552
1016 1545
1439 1892
1119 1541
984
1514
1043 1538
1023 1573
1022 1557
1032 1530
993
1580
964
1450
943
1482
1137 1568
1028 1532
1095 1580
RA401
Meta Sediments Ashland Wedowee Belt
0.11484
0.512073 (10)
RA 401(rerun) Meta Sediments Ashland Wedowee Belt
0.11510
0.512071 (12)
RA1005
Meta Sediments Ashland Wedowee Belt
0.11987
0.512124 (14)
RA 1021
Meta Sediments Ashland Wedowee Belt
0.12246
0.511934 (10)
RA 1029
Meta Sediments Ashland Wedowee Belt
0.10005
0.511926 (09)
RA 422A
Meta Sediments Ashland Wedowee Belt
0.11864
0.512132 (09)
RA536
Meta Sediments Ashland Wedowee Belt
0.11418
0.512071 (10)
RA 550
Meta Sediments Ashland Wedowee Belt
0.12368
0.512146 (14)
RA 587
Meta Sediments Ashland Wedowee Belt
0.12098
0.512128 (14)
RA701b
Meta Sediments Ashland Wedowee Belt
0.11443
0.512079 (16)
RA1034
Meta Sediments Ashland Wedowee Belt
0.12853
0.512192 (10)
Sm639
Meta Sediments Ashland Wedowee Belt
0.10907
0.512082 (10)
SC25
Canton Schist
PCF
0.10768
0.512085 (10)
Sc33
Canton Schist
PCF
0.10330
0.511939 (10)
NG057
Canton Schist
PCF
0.11546
0.512088 (11)
DR136
Canton Schist
PCF
0.11422
0.512043 (12)
PCF - Pumpkinvine Creek Amphibolites.
2σ errors on 143Nd/144 Nd ratios reported as last two significant digits. 143Nd/144 Nd measured ratio normalized to
146
Nd/144 Nd = 0.7219.
TDM ages calculated as in Salters and Stracke (2004).
TCHUR were calculated assuming the present day 143Nd/144NdCHUR = 0.512638 and 147Sm/144NdCHUR = 0.1967.
30
Table 1.7. U-Pb analysis of Galts Ferry Gneiss zircons by LA-MS-ICPMS
CORRECTED CONCENTRATIONS AND RATIOS
CALCULATED AGES & 1-S SD RANDOM ERRORS
207
206
206
U (ppm)
Pb
Pb*
± (%)
Pb* ± (%) error 206Pb* ± (Ma) 207Pb* ± (Ma) 206Pb* ± (Ma)
204
Pb
219
375
64
210
226
347
170
627
475
317
447
326
285
794
954
148
313
237
173
593
341
183
404
681
337
282
270
190
610
206
285
246
419
41
53
116
173
231
35
76
84
60
127
8718
9917
5578
3833
9485
13517
13614
7768
4088
7301
22930
3943
5351
3773
25160
5310
10278
14475
14257
2779
14965
19486
9046
6720
7041
9301
11697
14622
8184
11679
2451
14898
2549
2246
6566
5384
7466
12157
1501
9176
5774
3336
6848
235
238
U
0.5732
0.5507
0.5768
0.6532
0.5747
0.5577
0.6312
0.5002
0.4935
0.5901
0.6592
0.5317
0.5632
0.5355
0.5888
0.5606
0.5405
0.5343
0.5878
0.4710
0.6273
0.5790
0.4888
0.4406
0.5473
0.5256
0.5628
0.6058
0.5307
0.5253
0.5220
0.5714
0.3914
0.7638
0.6233
0.5919
0.5970
0.5781
0.7258
0.5814
0.6223
0.7106
0.5567
U
4.64
3.21
13.74
6.70
6.77
3.93
5
5.42
6.26
2.54
2.84
6.86
3.77
1.83
3.02
7.35
5.62
2.60
6.44
9.20
4.79
4.64
2.16
2.24
5.84
5.34
6.71
4.16
3.74
4.25
11.23
3
10.69
19.03
18.43
14.38
5.73
5.70
21.97
14.04
16.85
11.25
10.71
0.0737
0.0705
0.0698
0.0820
0.0759
0.0728
0.0777
0.0652
0.0638
0.0742
0.0825
0.0723
0.0733
0.0659
0.0734
0.0765
0.0695
0.0694
0.0771
0.0612
0.0819
0.0743
0.0637
0.0581
0.0721
0.0667
0.0720
0.0775
0.0672
0.0720
0.0742
0.0727
0.0485
0.0817
0.0734
0.0744
0.0785
0.0750
0.0767
0.0746
0.0741
0.0741
0.0732
corr.
0.70
1.66
3.80
4.27
0.72
1.87
1.37
5.21
3.67
1.47
2.07
2.92
1.52
1.39
2.59
1.29
5.06
1.17
2.01
4.82
2.25
0.75
0.87
1.79
4.52
2.21
6.02
0.87
3.37
0.71
1.08
1.83
7.83
5.78
1.57
1.21
4.47
0.82
4.54
4.24
1.86
5.14
1.42
31
0.15
0.52
0.28
0.85
0.11
0.48
0.27
0.96
0.59
0.58
0.73
0.43
0.40
0.76
0.86
0.18
0.90
0.45
0.31
0.52
0.47
0.16
0.40
0.80
0.77
0.41
0.90
0.21
0.90
0.17
0.10
0.61
0.73
0.30
0.09
0.08
0.78
0.14
0.21
0.30
0.11
0.46
0.13
238
235
U
458.5
439.3
435.5
508.2
471.9
453.3
482.7
407.6
399.2
462.6
511.4
450.2
457.0
411.5
457.1
475.6
433.3
433.1
478.8
383.2
505.8
463.2
398.3
364.6
449.2
416.7
448.4
481.5
419.4
448.5
461.9
452.6
305.4
506.5
456.9
462.8
487.5
466.4
476.8
464.3
461.0
460.9
455.4
207
U
3.1
7.1
16.0
9.7
3.3
8.2
6.4
20.6
14.2
5.2
10.2
12.7
6.7
5.5
11.4
5.9
21.2
4.9
9.3
17.9
6.0
6.3
3.4
6.3
19.6
8.9
26.1
4.0
13.7
3.1
4.8
8.0
23.4
28.2
6.9
5.4
21.0
3.7
20.9
19.0
8.3
22.9
6.3
460.1
445.5
462.5
510.4
461.1
450.0
496.9
411.8
407.3
470.9
514.1
432.9
453.6
435.5
470.1
451.9
438.8
434.7
469.5
391.9
494.4
463.9
404.1
370.7
443.2
428.9
453.4
481.0
432.3
428.7
426.5
458.9
335.4
576.2
491.9
472.1
475.3
463.2
554.1
465.4
491.3
545.1
449.4
Pb*
17.2
11.6
51.1
17.1
25.1
14.3
19.7
18.3
21.0
9.6
11.5
24.2
13.8
6.5
11.4
26.8
20.0
9.2
24.2
29.9
18.8
17.3
7.2
6.9
21.0
18.7
24.6
15.9
13.2
14.9
39.1
11.1
30.5
83.9
72.0
54.3
21.8
21.2
94.1
52.5
65.7
47.5
38.9
468
477
599
521
408
433
563
436
454
517
526
342
441
564
534
333
467
443
424
443
433
472
438
409
412
495
478
478
502
324
240
491
549
862
658
517
417
448
886
471
635
915
419
102
61
287
90
151
77
105
33
113
46
43
141
77
26
34
164
54
52
137
174
94
101
44
30
83
107
66
90
36
95
258
53
159
379
397
316
80
126
449
297
363
206
238
Table 1.7 continued
CORRECTED CONCENTRATIONS AND RATIOS
CALCULATED AGES & 1-S SD RANDOM ERRORS
207
206
206
206
U (ppm)
Pb
Pb*
± (%)
Pb* ± (%) error
Pb* ± (Ma) 207Pb* ± (Ma) 206Pb* ± (Ma)
204
Pb
164
77
53
39
76
46
174
72
527
5656
3936
3215
1646
6954
3292
5871
6875
20054
235
238
U
0.5476
0.5948
0.5757
0.7062
0.6275
0.7175
0.5549
0.6262
0.5720
U
9.79
7.63
21.86
22.29
16.37
10.24
9.70
18.31
1.90
0.0749
0.0735
0.0751
0.0767
0.0748
0.0741
0.0676
0.0729
0.0737
corr.
2.33
3.96
3.33
2.43
4.85
1.58
1.37
4.71
0.70
0.24
0.52
0.15
0.11
0.30
0.15
0.14
0.26
0.37
238
235
U
466.0
457.4
467.1
476.6
465.2
460.9
422.2
453.9
458.9
207
U
10.5
17.5
15.0
11.1
21.8
7.0
5.6
20.6
3.1
443.4
474.0
461.7
542.5
494.5
549.2
448.2
493.8
459.3
Pb*
35.2
28.9
81.3
93.9
64.2
43.5
35.2
71.7
7.0
328
555
435
830
633
935
584
683
462
216
142
486
467
338
208
209
381
39
U concentration has an uncertainty of ~15%.
Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma.
Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for
207
Pb/204Pb.
32
Table 1.8. U-Pb analysis of detrital zircons from meta-sandstone by LA-MS-ICPMS
Isotopic ratios
U
206
(ppm)
204
215
138
174
173
822
77
57
37
58
283
145
169
75
66
134
68
201
84
507
144
407
334
240
162
582
126
39
194
299
321
336
363
278
298
438
734
32
178
422
70
193
1072
351
Pb
Pb
16371
14899
6931
17202
55166
8342
5443
4455
6783
5968
9405
18110
11913
7035
10834
7642
16751
4102
40410
18133
41030
33630
22274
19728
54839
9808
2046
6197
35883
28831
46426
13379
24956
25680
59290
52263
4568
28587
25585
3864
31361
73111
79746
207
Pb*
± (%)
235
206
Pb*
± (%) error
238
U
2.028
1.875
1.653
2.121
1.753
1.786
2.123
1.792
1.749
0.666
1.681
3.637
2.205
1.849
2.273
2.539
2.067
1.901
1.860
1.749
2.801
1.790
1.960
1.727
2.078
0.808
1.559
1.652
1.817
2.408
2.168
1.743
1.778
1.729
1.555
1.753
1.994
3.132
1.640
2.559
2.112
1.938
1.687
Apparent ages (Ma)
U
2.03
4.86
4.79
3.72
0.99
12.49
7.04
12.77
6.16
8.06
4.30
2.29
4.31
7.92
2.59
5.25
3.10
9.70
1.66
4.79
1.57
1.73
1.77
4.10
1.16
9.96
18.57
5.46
2.85
2.42
2.47
2.86
3.10
2.54
5.87
1.22
16.41
3.19
5.70
9.15
4.11
1.57
5.33
0.1934
0.1752
0.1713
0.1970
0.1706
0.1683
0.1922
0.1780
0.1729
0.0819
0.1691
0.2773
0.2011
0.1707
0.2122
0.2216
0.1934
0.1761
0.1796
0.1717
0.2336
0.1772
0.1879
0.1741
0.1906
0.0955
0.1493
0.1712
0.1807
0.2120
0.1970
0.1721
0.1755
0.1717
0.1555
0.1686
0.1757
0.2471
0.1597
0.2134
0.1912
0.1797
0.1649
corr.
0.90
0.70
0.92
1.20
0.70
3.70
1.10
2.30
1.10
4.40
1.20
1.10
1.20
1.70
1.10
0.91
1.30
6.80
0.70
1.60
0.70
1.40
0.90
1.70
0.70
0.90
10.10
2.91
0.90
1
1.10
0.90
0.70
0.80
5.70
0.70
2.80
2.20
5.10
5.10
1.30
1.40
4.40
0.44
0.14
0.19
0.32
0.71
0.30
0.16
0.18
0.18
0.55
0.28
0.48
0.28
0.21
0.42
0.17
0.42
0.70
0.42
0.33
0.45
0.81
0.51
0.41
0.61
0.09
0.54
0.53
0.32
0.41
0.45
0.32
0.23
0.32
0.97
0.57
0.17
0.69
0.90
0.56
0.32
0.89
0.83
33
206
Pb*
± (Ma)
238
Pb*
± (Ma)
235
U
1140.2
1040.8
1019.4
1159.4
1015.6
1002.8
1133.6
1056.1
1028.4
507.5
1007.4
1578.2
1181.6
1016.2
1241.0
1290.6
1139.8
1045.8
1064.9
1021.6
1353.8
1051.9
1110.0
1035.1
1125.0
588.0
897.3
1019.2
1071.0
1239.8
1159.5
1024.0
1042.6
1021.8
931.9
1004.7
1043.7
1423.7
955.6
1247.1
1128.0
1065.7
984.2
207
1125.2
1072.4
990.8
1155.8
1028.4
1040.4
1156.4
1042.6
1027.1
518.7
1001.7
1557.8
1182.6
1063.4
1203.9
1283.5
1138.2
1081.5
1067.2
1026.9
1355.9
1042.0
1102.1
1018.7
1141.7
601.6
954.2
990.4
1051.6
1245.1
1171.1
1024.6
1037.7
1019.5
952.7
1028.6
1113.5
1440.6
985.7
1289.0
1152.9
1094.4
1003.8
Pb*
± (Ma)
207
U
9.4
6.7
8.7
12.7
6.6
34.4
11.5
22.4
10.5
21.5
11.2
15.4
13.0
16.0
12.4
10.6
13.6
65.7
6.9
15.1
8.5
13.6
9.2
16.3
7.2
5.1
84.6
27.4
8.9
11.3
11.7
8.5
6.7
7.6
49.5
6.5
27.0
28.1
45.3
57.8
13.5
13.8
40.2
206
Pb*
13.8
32.2
30.3
25.7
6.4
81.5
48.6
83.4
39.8
32.7
27.4
18.2
30.1
52.2
18.3
38.3
21.2
64.6
11.0
31.0
11.7
11.2
11.9
26.4
7.9
45.2
115.4
34.5
18.7
17.4
17.2
18.4
20.1
16.3
36.3
7.9
111.4
24.5
35.9
66.9
28.4
10.5
34.0
1096
1137
928
1149
1056
1120
1199
1014
1024
568
989
1530
1184
1162
1138
1272
1135
1154
1072
1038
1359
1021
1086
984
1173
653
1088
927
1012
1254
1193
1026
1027
1014
1001
1080
1252
1466
1053
1360
1200
1152
1047
36
96
97
70
14
239
137
256
123
147
84
38
82
153
47
101
56
137
30
91
27
20
30
76
18
213
314
95
55
43
44
55
61
49
29
20
318
44
51
147
77
14
61
Table 1.8 continued
Isotopic ratios
U
206
(ppm)
204
213
214
35
316
385
277
69
160
202
68
112
213
422
261
521
126
542
52
137
421
246
344
496
170
173
141
238
319
282
501
169
395
90
755
881
110
500
112
75
100
Pb
Pb
10787
18812
5637
28898
26025
11477
6065
16561
10405
3995
10297
21588
13076
7741
5602
10457
49461
5254
3599
14335
12829
29772
74843
18221
21924
16160
29975
43306
44281
30093
16376
14769
8141
25311
53077
4675
63345
17631
7521
8509
207
Pb*
± (%)
235
206
Pb*
± (%)
238
U
26
2.239
2.203
2.972
1.703
1.719
1.805
1.608
1.539
1.680
1.692
1.666
2.266
1.994
1.571
1.511
1.764
2.942
2.413
2.242
27
2.302
1.735
1.663
2.661
2.142
2.533
1.726
2.142
2.037
1.957
1.579
1.580
2.873
1.679
1.547
1.874
2.146
2.040
1.427
Apparent ages (Ma)
U
3.55
2.30
9.51
3.45
2.23
3.86
7.46
6.56
7.46
7.74
8.76
2.70
2.23
4.62
3.97
6.52
1.15
7.68
6.73
2.13
8.70
1.36
1.66
6.17
2.41
5.11
2.09
2.37
2.22
1.40
2.71
1.95
5.72
1.29
1.53
6.17
1.72
4.65
6.36
9.34
0.1894
0.2069
0.1935
0.2322
0.1691
0.1692
0.1675
0.1637
0.1543
0.1590
0.1669
0.1708
0.1903
0.1861
0.1526
0.1528
0.1702
0.2496
0.2154
0.2006
0.1814
0.2020
0.1722
0.1705
0.2235
0.2017
0.2205
0.1671
0.2015
0.1851
0.1801
0.1534
0.1583
0.2270
0.1590
0.1475
0.1786
0.1985
0.1884
0.1400
error
corr.
2
1.10
3.40
3
1.80
2.30
1.30
1.80
6.20
1.51
1.90
0.90
1.80
2.51
3.70
1.90
0.70
0.71
2.34
1.70
8.30
0.70
0.70
4.10
1.70
1.90
1.20
1.40
0.70
1.20
1.90
0.70
2.50
0.80
1.30
2.70
1
2.50
1.20
5.20
0.56
0.48
0.36
0.87
0.81
0.60
0.17
0.27
0.83
0.19
0.22
0.33
0.81
0.54
0.93
0.29
0.61
0.09
0.35
0.80
0.95
0.52
0.42
0.66
0.70
0.37
0.57
0.59
0.32
0.86
0.70
0.36
0.44
0.62
0.85
0.44
0.58
0.54
0.19
0.56
206
Pb*
± (Ma)
238
Pb*
± (Ma)
235
U
1118.6
1212.3
1140.6
1346.4
1007.4
1007.9
998.5
977.8
925.3
951.4
995.1
1016.8
1123.3
1100.6
915.6
916.8
1013.6
1436.5
1257.8
1179.0
1074.8
1186.3
1024.5
1015.4
1300.8
1184.8
1284.8
996.1
1183.6
1095.3
1068.0
920.5
947.3
1318.8
951.2
887.3
1059.4
1167.7
1112.8
845.1
207
1117.7
1193.5
1182.0
1400.6
1010.0
1015.8
1047.3
973.5
946.4
1001.1
1005.6
995.7
1202.0
1113.5
958.9
935.1
1032.6
1392.8
1246.7
1194.3
1118.0
1213.0
1021.9
994.7
1317.8
1162.6
1281.6
1018.4
1162.6
1128.0
1100.9
962.3
962.6
1375.1
1000.8
949.5
1072.2
1163.9
1129.3
900.3
Pb*
± (Ma)
207
U
20.6
12.2
35.6
36.5
16.8
21.5
12.0
16.3
53.4
13.3
17.5
8.5
18.6
25.4
31.6
16.2
6.6
9.1
26.7
18.3
82.2
7.6
6.6
38.5
20.0
20.6
14.0
12.9
7.6
12.1
18.7
6.0
22.0
9.5
11.5
22.4
9.8
26.7
12.3
41.2
206
Pb*
24.1
16.2
66.5
26.2
14.3
24.8
48.8
41.1
46.0
49.3
55.9
17.1
15.7
31.3
24.6
39.9
7.4
58.3
48.3
14.9
59.1
9.6
10.7
39.2
17.8
35.4
15.2
15.3
15.4
9.6
18.2
12.1
35.6
9.7
9.7
38.1
11.4
32.2
43.4
55.8
1116
1160
1259
1484
1015
1033
1150
964
996
1111
1028
949
1346
1139
1059
978
1073
1326
1228
1222
1203
1261
1016
949
1346
1121
1276
1067
1124
1192
1166
1059
998
1464
1111
1097
1098
1157
1161
1038
59
40
174
32
27
63
146
129
84
152
173
52
26
77
29
127
18
148
124
25
52
23
30
95
33
95
33
39
42
14
38
37
105
19
16
111
28
78
124
157
U concentration has an uncertainty of ~15%.
Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma.
Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for
207
Pb/204Pb.
34
Table 1.9. U-Pb analysis of Mulberry Rock Gneiss zircons by LA-MS-ICPMS
CORRECTED CONCENTRATIONS AND RATIOS
U (ppm)
206
Pb
204
Pb
245
283
444
149
661
515
309
558
218
1625
920
375
296
794
794
823
1776
1024
1409
375
793
316
287
842
379
459
962
264
1005
147
648
273
248
285
390
672
316
980
272
805
262
1204
224
464
1079
574
836
1486
4205
2680
4416
1194
3255
1692
1394
6381
3075
3898
8917
2248
1637
6017
2804
23984
1844
2165
3143
1183
5679
1694
1970
6576
2027
1735
1609
2484
1176
1056
3007
4770
6596
1087
2381
2701
2116
2674
3246
1813
3850
654
2528
1028
207
Pb*
± (%)
235
Pb*
U
6.56
10.54
6.60
8.30
4.06
13.32
7.08
8.06
8.43
1.81
4.26
4.58
6.32
5.75
7.08
4.37
2.17
1.43
7.57
5.29
4.72
11
3.89
5.52
8.38
3.81
4.69
7.21
3.26
10.96
6.03
10.85
8.38
8.27
4.95
4.60
8.25
4.40
7.35
7.47
5.33
6.45
8.46
15.99
5.89
6.24
0.0757
0.0748
0.0707
0.0720
0.0668
0.0491
0.0678
0.0658
0.0689
0.0709
0.0666
0.0710
0.0698
0.0674
0.0675
0.0684
0.0578
0.1707
0.0675
0.0689
0.0618
0.0503
0.0713
0.0672
0.0673
0.0683
0.0630
0.0729
0.0703
0.0727
0.0625
0.0762
0.0707
0.0686
0.0684
0.0655
0.0717
0.0664
0.0619
0.0680
0.0770
0.0685
0.0764
0.0613
0.0718
0.0498
CALCULATED AGES & 1-S SD RANDOM ERRORS
± (%) error
238
U
0.6710
0.5953
0.5569
0.5949
0.5370
0.4206
0.5412
0.5439
0.5681
0.5627
0.5398
0.5712
0.5501
0.5507
0.5560
0.5322
0.4522
1.7656
0.5576
0.5283
0.4868
0.4192
0.5384
0.5533
0.5616
0.5170
0.5101
0.6279
0.5761
0.6074
0.5045
0.6552
0.6108
0.5001
0.5360
0.5379
0.5829
0.5389
0.4984
0.5737
0.6107
0.5612
0.5598
0.5680
0.5916
0.3923
206
corr.
4.16
1.69
1
2.74
1.86
8.38
1.39
3.31
2.11
1.34
3.29
1
1.14
1.24
3.25
1.77
1.67
1
2.01
2.46
2.68
6.12
1.01
1.22
1.94
1.86
2.96
2.79
2.06
1.82
2.94
1.49
2.83
2.14
1.28
1.93
2.53
2.90
2.27
2.91
1.82
2.15
3.92
2.19
2.26
2.56
0.64
0.16
0.15
0.33
0.46
0.63
0.20
0.41
0.25
0.74
0.77
0.22
0.18
0.22
0.46
0.41
0.77
0.70
0.26
0.47
0.57
0.56
0.26
0.22
0.23
0.49
0.63
0.39
0.63
0.17
0.49
0.14
0.34
0.26
0.26
0.42
0.31
0.66
0.31
0.39
0.34
0.33
0.46
0.14
0.38
0.41
35
206
Pb*
± (Ma)
238
Pb*
± (Ma)
235
U
470.8
465.3
440.6
448.4
417.3
309.3
423.3
410.9
429.9
441.9
415.7
442.5
435.3
420.7
421.5
426.6
362.3
1016.2
421.1
430.0
386.9
316.8
444.3
419.3
419.9
426.3
394.0
453.6
438.4
452.8
391.2
473.4
440.6
427.7
427.0
409.1
446.6
414.6
387.7
424.1
478.6
427.3
474.7
383.7
447.3
313.3
207
521.4
474.3
449.5
474.0
436.5
356.5
439.3
441.0
456.8
453.3
438.3
458.8
445.1
445.5
449.0
433.3
378.9
1032.9
450.0
430.7
402.7
355.5
437.4
447.2
452.6
423.2
418.5
494.8
462.0
481.9
414.8
511.7
484.1
411.8
435.8
437.0
466.3
437.7
410.6
460.4
484.0
452.3
451.4
456.8
471.9
336.1
Pb* ± (Ma)
207
U
18.9
7.6
4.3
11.9
7.5
25.3
5.7
13.2
8.8
5.7
13.2
4.3
4.8
5.0
13.3
7.3
5.9
9.4
8.2
10.3
10.1
18.9
4.3
4.9
7.9
7.7
11.3
12.2
8.7
8.0
11.1
6.8
12.1
8.9
5.3
7.7
10.9
11.7
8.6
12.0
8.4
8.9
18.0
8.2
9.8
7.8
206
Pb*
26.7
39.9
24.0
31.4
14.4
40.1
25.3
28.9
31.0
6.6
15.2
16.9
22.8
20.7
25.7
15.4
6.9
9.2
27.5
18.6
15.7
33.0
13.8
20.0
30.6
13.2
16.1
28.2
12.1
42.1
20.5
43.6
32.3
28.0
17.5
16.3
30.9
15.6
24.8
27.7
20.5
23.6
30.8
58.9
22.2
17.9
749
518
496
600
539
676
524
601
595
512
559
541
496
576
592
469
482
1068
601
435
494
617
401
593
622
406
556
690
581
623
548
687
696
323
483
587
565
561
542
646
510
582
335
843
593
496
107
229
144
170
79
222
152
159
177
27
59
98
137
122
137
89
31
20
158
104
86
198
84
117
176
74
79
142
55
234
115
230
168
182
106
91
171
72
153
148
110
132
170
331
118
126
Table 1.9 continued
U (ppm)
CORRECTED CONCENTRATIONS AND RATIOS
207
206
206
Pb
Pb*
± (%)
Pb*
± (%)
204
235
Pb
1151
311
269
165
269
566
298
169
705
299
176
664
158
236
1904
1362
550
296
364
187
121
238
U
1131
168
430
963
2564
875
610
576
620
1551
5945
6502
5733
2538
3008
1343
1414
6614
755
10115
2230
U
0.4870
0.6287
0.5408
0.5382
0.4156
0.5200
0.5449
0.9735
0.6253
0.5337
0.5503
0.4763
0.6770
0.6055
0.5652
0.6211
1.2266
0.5906
0.7355
0.6460
0.5724
3.41
9.17
9.92
13.39
15.17
6.95
12.19
10.42
7.20
7.29
9.50
9.13
11.09
10.17
3.97
2.96
4.26
5.15
7.47
7.66
11.61
0.0589
0.0532
0.0545
0.0653
0.0498
0.0633
0.0606
0.0954
0.0724
0.0641
0.0711
0.0645
0.0878
0.0759
0.0698
0.0750
0.1089
0.0750
0.0724
0.0784
0.0770
error
CALCULATED AGES & 1-S SD RANDOM ERRORS
207
206
Pb*
± (Ma)
Pb*
± (Ma)
Pb*
± (Ma)
206
238
corr.
2.47
1.89
3.25
5.34
6.84
4.59
4.75
5.48
4.07
2.81
2.06
7.63
1.12
2.64
2.44
1.21
3.01
0.95
3.65
4.19
0.93
0.73
0.21
0.33
0.40
0.45
0.66
0.39
0.53
0.57
0.39
0.22
0.84
0.10
0.26
0.61
0.41
0.71
0.19
0.49
0.55
0.08
235
U
207
U
368.9
334.6
342.3
408.1
313.3
396.2
379.3
587.4
450.6
400.8
443.3
403.0
542.7
471.7
435.3
466.6
666.4
466.6
451.1
486.8
478.2
8.9
6.2
10.9
21.1
20.9
17.6
17.5
30.8
17.7
10.9
8.8
29.8
5.8
12.0
10.3
5.4
19.1
4.3
15.9
19.6
4.3
402.9
495.3
439.0
437.2
352.9
425.2
441.7
690.3
493.2
434.3
445.2
395.5
525.0
480.7
454.9
490.6
812.8
471.3
559.8
506.0
459.6
Pb*
11.3
36.0
35.4
47.6
45.3
24.1
43.7
52.2
28.1
25.8
34.2
29.9
45.5
39.0
14.6
11.5
23.9
19.4
32.2
30.5
42.9
603
582
984
594
622
585
781
720
696
616
455
352
449
524
555
604
51
97
191
267
293
113
237
192
127
145
206
113
246
216
68
59
494
620
594
368
112
134
139
261
U concentration has an uncertainty of ~15%.
Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma.
Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for
207
Pb/204Pb.
Table 1.10 REE concentrations of the metasediments of the eastern Blue Ridge.
RA401 RA 401 RA701b RA1034 Sc33
NG057 DR136 RA 422A RA536 RA 550 RA 587 RA1005 RA 1029 RA 1021 Sm639
SC25
La
36
31
73
30
203
38
56
27
33
20
35
65
73
18
120
144
Ce
55
36
180
84
327
93
127
35
71
45
74
88
95
58
208
132
Pr
5.4
5.1
13.5
5.0
30.3
10.8
10.5
7.6
8.5
5.6
9.5
11.0
12.0
3.4
25
23
Nd
20
19
48
19
102
40
41
29
32
22
37
41
44
13
88
78
Sm
3.7
3.6
8.9
3.9
17.0
7.5
7.6
5.6
6.0
4.3
7.2
8.1
7.9
2.5
16
14
Eu
0.69
0.76
0.88
0.64
1.65
1.28
1.11
1.18
1.02
1.01
1.94
1.66
1.39
0.51
1.81
1.65
Gd
1.20
2.08
3.14
1.46
5.49
2.64
2.65
3.03
3.25
2.59
4.09
4.93
4.59
1.72
4.43
3.16
Tb
0.26
0.35
0.60
0.34
1.25
0.50
0.51
0.55
0.57
0.55
0.75
1.00
0.85
0.33
1.06
1.00
Dy
1.90
1.85
6.25
5.38
9.89
4.14
3.73
3.11
3.29
3.99
4.39
6.47
5.60
2.47
10.87
8.57
Ho
0.32
0.30
1.29
1.16
1.92
0.82
0.74
0.49
0.58
0.83
0.72
1.21
1.11
0.47
2.10
1.55
Er
0.81
0.77
4.10
3.55
6.16
2.49
2.27
1.16
1.75
2.54
1.71
3.63
3.28
1.34
6.18
4.43
Yb
0.69
0.72
3.20
2.49
6.37
2.25
2.04
0.94
2.34
2.80
1.23
4.64
3.31
1.37
4.26
3.57
Lu
0.12
0.13
0.58
0.46
1.23
0.40
0.36
0.15
0.43
0.47
0.18
0.82
0.54
0.23
0.68
0.61
36
Study area
(Fig 1.1 A)
Figure.1.1. Generalized geological map of the Southern Appalachian. CBR – Central Blue
Ridge; DGB – Dahlonega Gold Belt. Modified after Holm & Das, 2006.
37
Figure 1.1A. Generalized geologic map of eastern Alabama and western Georia Blue Ridge terranes showing
Fig1.1A
the tectonic relationship between the PCG and HG. Modified after Holm and Das, 2006 (in review). WBRwestern Blue Ridge; PCF-Pumpkinvine Creek Formation; MRG-Mulberry Rock Gneiss; AG-Austell Gneiss;
HG-Hillabee Greenstone; EQD-Elkahatchee Quartz Diorite.
38
Figure 1.2. Geochemical representation of bimodality found in the HG and Hillabee
dacite and PCF amphibolites and felsic gneiss (GFG). (A) SiO2 weight % histogram of
the felsic and mafic components of the HG belt. (B) SiO2 weight % histogram of the
felsic and mafic components of the PCF. Data from this study, Durham, 1993 and
McConnell and Abrams, 1984.
39
Figure 1.3. Total Alkalis versus SiO2 (Lebas et al. 1986) place the mafic component
of the PCF (purple filled diamond) and HG (blue filled square) in the basalt field
while the GFG (purple open diamond) falls in the rhyolite field and the Hillabee
dacite (blue open square) falls in the dacite field. Data from Durham, 1993 and
McConnell and Abrams, 1984 are plotted too. Literature data symbols: GFG (pink
open plus sign), Hillabee dacite (blue open multiply sign), PVC amphibolite (pink
filled plus sign), greenstones (blue open multiply sign).
40
Figure 1.4A. Classification of the Hillabee dacite and GFG using Shand’s index of
Maniar and Piccoli (1989). Both the felsic components display a peraluminous
affinity. Symbols and data source same as Figure 1.3.
Figure 1.4B. AFM diagram after Irvine and Barager (1971) showing tholeiitic
trend of HG and PCF amphibolites and calc-alkaline trend of Hillabee dacite and
GFG. Symbols same as Figure 1.3. FeO* is the total FeO.
41
A
Ti (ppm)
100000
10000
Expected
compositio
n of felsic
component
if derived
from
fractional
crystallizat
ion of
mafic
1000
100
10
100
1000
Zr (ppm)
100000
Ti (ppm)
B
10000
Expected
composition
of felsic
component if
derived from
fractional
crystallizatio
n of mafic
material.
1000
100
10
100
1000
Zr (ppm)
Figure.1.5. Ti vs. Zr plot to illustrate trends of A) HG and Hillabee dacite B)PCF
amphibolites and GFG and expected trend of dacite and GFG if they were derived by
fractional crystallization of their mafic counterparts. Symbols same as Figure. 1.3.
42
Ti (ppm)
100000
10000
1000
10
60
110
160
210
Zr (ppm)
10
Nb (ppm)
9
8
7
6
5
4
3
2
1
0
10
60
110
160
210
160
210
Zr (ppm)
60
Y (ppm)
50
40
30
20
10
0
10
60
110
Zr (ppm)
Figure.1.6. Co-variation diagrams of relatively immobile
elements for the HG (blue filled squares) and PCF
amphibolites (pink filled diamonds). Ti, Nb and Y vs. the
HFSE Zr.
43
A
B
Figure 1.7. Tectonic discrimination diagrams for the PCF amphibolites
and HG. (A) Zr/4-Nb*2-Y diagram (Meschede, 1986). (B) Y vs. Nb
diagram (Pearce et al., 1984). OIB- ocean island basalt, WP Alk- withinplate alkaline; WP Th- within-plate tholeiite; E-MORB- enriched midocean ridge basalt; N-MORB- normal mid-ocean ridge basalt; VABvolcanic arc basalt. Symbols same as Figure 1.6.
44
Figure1.8. La/10-Y/15 Nb/8 diagram (Cabanis and Lecolle, 1989). IAT- island-arc
tholeiite, OFB- ocean-floor basalt, MORB- mid-ocean ridge basalt, BABB- back-arc
basin basalt, VAT- volcanic-arc tholeiite, NMORB- normal mid-ocean ridge basalt,
EMORB- enriched mid-ocean ridge basalt, Cont. Continental. Symbols same as Figure
1.6.
45
Figure 1.9. Granitoid tectonic discrimination diagrams of GFG (pink open
diamonds) and Hillabee dacite (blue open squares). (A) Y vs. Nb diagram
(Pearce et al., 1984). (B) Y + Nb vs. Rb diagram (Pearce et al., 1984). VAGvolcanic arc granite, syn-COLG- syn-collisional granite, WPG- within plate
granite, ORG- ocean-ridge granite.
46
Rock/Primitive Mantle
McDonough and Sun
100.00
PCF
10.00
1.00
0.10
0.01
Rb Ba Th
U
Nb La Ce Pb Pr
Sr Nd
Zr Sm Eu
Ti
Dy Yb
Y
Lu
100.00
Hillabee Greenstone
10.00
1.00
0.10
0.01
Rb Ba
Th
U
Nb
La
Ce Pb
Pr
Sr
Nd
Zr Sm Eu
Ti
Dy
Yb
Y
Lu
Figure. 1.10. Primitive Mantle- normalized (McDonough and Sun, 1995) spider
diagram plot of PCF amphibolites and Hillabee Greenstone.
47
Rock/Chondrites
McDonough and Sun 1995
100
10
PCF Amphibolites
1
La
Ce
Pr
Nd
Sm
Eu
Tb
Dy
Ho
Er
Yb
Lu
100
Hillabee Greenstone
10
1
La
Ce
Pr
Nd
Sm
Eu
Tb
Dy
Ho
Er
Yb
Lu
Figure. 1.11. CI chondrite-normalized (McDonough and Sun, 1995) REE
diagram plotting of PCF amphibolites and Hillabee Greenstone.
48
Rock/Chondrites
McDonough and Sun 1995
1000
GFG
100
10
1
La
Ce
Pr
Nd Sm Eu Tb
Dy Ho
Er
Yb
Lu
Yb
Lu
1000
Hillabee dacite
100
10
1
La
Ce
Pr
Nd Sm
Eu
Tb
Dy
Ho
Er
Figure. 1.12. CI chondrite-normalized (McDonough and Sun, 1995) REE diagram of
GFG and Hillabee dacites. Field for metadacite from Durham, 1993 is shown for
comparison.
49
9.00
7.00
HG
5.00
3.00
εNd
Hillabee
Metadacite
1.00
GFG
-1.00
-3.00
PCF
Amphibolite
-5.00
-7.00
0.703
0.704
0.705
0.706
0.707
0.708
0.709
(87Sr/86Sr)i
Figure.1.13. Initial Sr vs. εNd isotopic plot of Pumpkinvine Creek Formation and
Hillabee Greenstone Sequence Rocks.
50
0.724
Y 131
391 Ma isochron
0.720
86
Sr/ Sr
NT 911
0.716
87
Y 24
Y 23 C
0.712
TA 7
BH 166
US 41A
Age = 391±110 Ma
0.708
Initial
87
86
Sr/ Sr =0.7076±0.0020
SC 318G
MSWD = 60
0.704
0.0
0.4
0.8
1.2
1.6
87
2.0
2.4
2.8
86
Rb/ Sr
0.720
NT 911
Figure. 1.14. Rb-Sr
isochron of the GFG
Y 23 C
0.712
BH 166
87
86
Sr/ Sr
0.716
TA 7
0.708
Age = 450±66 Ma
Initial
87
86
Sr/ Sr =0.7082±0.0012
MSWD = 2.2
0.704
0.0
0.2
0.4
0.6
0.8
87
1.0
1.2
1.4
86
Rb/ Sr
0.724
Y 131
0.716
86
Sr/ Sr
0.720
87
Y 24
0.712
US 41A
Age = 424±25 Ma
0.708
Initial
87
SC 318G
86
Sr/ Sr =0.7059±0.0010
MSWD = 0.8
0.704
0.0
0.4
0.8
1.2
1.6
87
86
Rb/ Sr
2.0
2.4
51
2.8
GFG, normalized to CI chondrite (McDonough & Sun 95)
250
High initial 87Sr/86Sr
sample/standard
200
150
100
50
Low initial 87Sr/86Sr
0
La Ce Pr Nd Sm Eu Gd Tb Dy Ho Er Yb Lu
Figure. 1.15. Chondrite normalized REE plot of GFG samples. The samples
define two groups. Light gray with high initial 87Sr/86Sr (~0.708) and dark gray
with low initial 87Sr/86Sr (~0.705).
52
data-point error ellipses are 2σ
0.11
650
550
450
0.07
206
Pb/238U
0.09
350
0.05
Concordia Diagram
Galts Ferry Gneiss Samples
250
0.03
0.2
0.4
0.6
207
Pb/
0.8
1.0
235
U
25
Number
20
15
10
5
0
260
300
340
380
420
460
500
540
206Pb/238U age in Ma
510
box heights are 1σ
206Pb/238U age
490
470
450
430
Mean = 460.4±9.7 95% conf.
MSWD = 1.7
410
53
Figure. 1.16. U-Pb ages of
zircons from GFG plotted in
Concordia diagram, histogram
and weighted average plot. Bin
width for the histogram is 20
Ma. 206Pb/238U ages of zircons
falling in the modal class and
second highest frequency class
after the modal class of the
histogram are selected for the
weighted average plot.
462+/-5 Ma
457+/-7 Ma
200 micron
463+/-9 Ma
505+/-6 Ma
200 micron
Figure. 16.1A. Galts Ferry Gneiss zircon showing magmatic zoning (top) and
older core (bottom).
54
0.5135
0.5130
0
143Nd/144Nd
0.5125
500
1000
1500
2000
CHUR
0.5120
143
Nd/144Nd
Depleted Mantle extracted at 3.4 Ga.
0.5115
0.5110
0.5105
Grenville basement at 700 Ma.
0.5100
Age in Ma
Figure. 1.17. Nd isotopic evolution of the eastern Blue Ridge. CHUR (Hamilton et al.,
1983) evolution curve in bold red and Depleted Mantle (Salters and Stracke, 2004)
evolution curve in bold blue. The two green evolution curves are for the two outliers
with the least TDM model age and the red evolution curve is for the outlier with the
oldest TDM model age. Nd isotopic value for the Grenville crust at 700 Ma is shown by
the gray shaded box (Carrigan et al., 2003; Hatcher et al., 2004).
55
NASC normalized REE plot of the Metasediments
5.0
4.5
4.0
3.5
3.0
2.5
2.0
1.5
1.0
0.5
0.0
La
Pr
Nd
Sm
Eu
Tb
Dy
Ho
Er
Yb
Lu
Figure.1.18. North American Shale normalized (Gromet, 1984) REE plot
of the metasediments from Ashland Wedowee belt and Canton Schist.
Two samples with youngest TDM is shown in green and the one oldest
sample is shown in red.
56
0.34
data-point error ellipses are 2σ
0.30
0.26
1400
206Pb/238U
0.22
0.18
1000
0.14
600
0.10
Concordia Diagram
Meta-sandstone from PCF
0.06
200
0.02
0
1
2
3
4
207Pb/235U
18
16
14
Number
12
10
8
6
4
2
0
300
500
700
900
1100
1300
1500
1700
207Pb-206Pb age in Ma
Figure. 1.19. Concordia Plot and Histogram (bin size = 50 Ma) of the zircons
from the meta-sandstone within the Pumpkinvine Creek Formation.
57
1.15
0.95
87
Sr/ 86Sr
1.05
0.85
Age = 467±16 Ma
0.75
Initial
87
86
Sr/ Sr =0.7076±0.0036
MSWD = 7
0.65
0
20
40
87
Rb/86Sr
Figure.1.20. Mulberry Rock Gneiss whole rock Rb-Sr
58
60
0.20
1100
data-point error ellipses are 2σ
0.16
206Pb/238U
900
0.12
700
500
0.08
300
0.04
Concordia Diagram
Mulberry Rock Gneiss
100
0.00
0.0
0.4
0.8
1.2
1.6
2.0
207Pb/235U
16
14
35 data in this cluster
12
Number
10
8
Lead Loss
6
Inhereted (?) or mixing
4
2
0
200
250
300
350
400
450
500
550
600
Ma
Figure 1.21. U-Pb ages of zircons from
Mulberry Rock Gneiss plotted in Concordia
diagram, and histogram.
59
Hillabee
dacite
Figure. 1.22. Epsilon Nd of Ordovician arc terrane felsic magmas of the
Appalachians. Crust fields and included arc terrane values are modified
from Coler et al. (2001) and references therein. The new data from the GFG
and Hillabee dacite are compared with the other more northern Appalachian
arc terranes and associated units
60
Figure. 1.23 Passive mechanisms of lithosphere extension. From a synthesis and
simplification of numerous studies that followed Celar Sengor and Burke, 1978. Modified
after Geoffroy, 2005.
Figure. 1.24. Accretionary Orogen in southern Appalachians (ca. Ord)
61
CHAPTER TWO
KILOMETERS SCALE STRONTIUM ISOTOPIC HOMOGENIZATION DURING
METAMORPHISM: A CASE STUDY IN THE TRES PEIDRAS GRANITE, NEW MEXICO
2.1 Introduction
More than a half century ago Compston and Jeffery (1959) made a major advance in the budding
science of geochronology when they showed that the common discordance between whole-rock
and mineral Rb-Sr dates of "granite-like" bodies were artifacts of the use of an initial strontium
isotopic composition in the calculation of dates. They revealed that, on the scale of a whole-rock
sample, a redistribution of Sr between minerals can occur during metamorphism, so that the whole
rock and minerals become homogenized to a common 87Sr/86Sr ratio. Furthermore it was indicated
that minerals and whole-rock yielded concordant dates for the time of metamorphism, when this
rehomogenized ratio was used as the common strontium in age calculations. Shortly after that
Nicolaysen (1961) of the Bernard Price Institute introduced a graphical solution to such systems.
Making use of the BPI plot (now referred to as isochron diagram), Nicolaysen demonstrated that
different granite-gneiss whole-rocks of the Baltimore Dome had remained closed systems during
metamorphism while there had been redistribution of Sr isotopes between the minerals of
individual rocks. Multiple whole rock samples yielded isochron ages that provided the time of
initial crystallization of the granite, while individual whole rocks and their respective minerals
yielded isochrons that date the time of metamorphic rehomogenization of strontium. The scale of
this strontium isotopic exchange was that of a whole-rock (decimeters).
There are numerous cases in the literature where these early advances in understanding have found
application. There are also numerous reports in which whole rocks appear to have been open Rb-Sr
systems, and analyses fail to produce the isochrons that should result as a necessary consequence of
close systems. Then there are reports in which well defined whole-rock isochrons yield dates less
than those obtained from U-Pb dates of zircons. Proposed reasons for such discrepancy between
well defined Rb-Sr whole rock isochron ages and zircon ages from the same rocks have been (1)
the presence of older zircon xenocrysts (Odom and Fullager, 1984), (2) a systematic loss of
radiogenic 87Sr or gain of 87Rb (Bickford and Mose 1975), (3) development of an apparent
isochron by mixing (Field and Raheim, 1980), and (4) an undefined "resetting" of an isochron age
by metamorphism (Page, 1978). It is inferred that such a resetting would involve an isotopic
rehomogenization of strontium on a scale equal to or greater than the distance between samples. At
present such large (kilometer) scale rehomogenization yet has been demonstrated. Now
geochronological investigations of the Tres Piedras Granite of Northcental New Mexico seem to
have revealed a remarkable example.
Tres Piedras Granite located near the town of Tres Piedras of the Rio Arriba County is one of the
several granites intruded during the growth of Proterozoic continental lithosphere in the
southwestern United States between 1.74 and 1.65 Ga (Maxon, 1976, Karlstrom and Bowring,
1988; Bauer and Williams, 1994). Two well documented thermal event is recorded in the area, at
ca. 1.6 Ga (Grambling and Dallmeyer, 1993) after the initial continental building period and at 1.48
Ga (Lanzirotti and Hanson, 1997) that affected the plutons in the area.
62
2.2 General Geology
Extensive Proterozoic continental lithosphere was formed from Wyoming southwestward till New
Mexico during Middle Proterozoic (~1800 – 1600 Ma) that record two major orogenic cycles; the
Yavapai Orogeny (~1700 Ma) and the Mazatzal Orogeny (~1650 Ma) (Hoffman, 1988; Karlstrom
and Bowring, 1988). Subsequently the rocks underwent regional and contact metamorphism and
intrusion of “anorogenic” granitoids at 1400 Ma (Karlstrom et al., 1997; Pedrick et al., 1998; Read
et al., 1999; Williams et al., 1999).
Controversies exist regarding metamorphic and deformation history of north central New Mexico.
While earlier studies (Grambling,1986; Bowring and Karlstrom, 1990; and Nyman et al., 1994)
propose one metamorphic event either at 1650 Ma or 1400 Ma, more recent works have identified
two deformation events at both 1650 Ma or 1400 Ma (Pedrick et al., 1998; Read et al., 1999;
Williams et al., 1999) In many Precambrian – cored uplifts of northern New Mexico and southern
Colorado, the ages of major fold and thrust structures, regional foliation and associated
metamorphism are loosely constrained between the ~1.70-1.65 Ga age of deformed plutons and the
~ 1.4Ga age of cross-cutting plutons, dykes and cooling of metamorphic minerals (Reed et al.,
1987; Gibson and Simpson, 1988; Williams, 1991). Assuming both the events were responsible for
the deformation and metamorphism of the rocks, it is still debated which event is responsible for
which deformational fabrics.
Proterozoic supracrustal rocks are exposed in isolated fault bounded uplifts across New Mexico.
Lithologies are dominated by metavolcanic rocks, metaquartzite and pelitic schist. Primary ages
range from 1750 to 1650 Ma, generally becoming younger to the south (Bowring and Condie,
1982). The supracrustal rocks were intruded by several granitic suites as well as several younger
batholiths prior to metamorphism. The oldest rocks of the Taos, Tusas and Santa Fe Ranges are
primarily gneisses and amphibolites. Here the Moppin, Pecos, and Gold Hill complexes have been
dated between 1760 to 1720 Ma (Bowring and Condie, 1982, Robertson and Condie, 1989,
Bowring, 1986, Silver and Dickinson 1987). Stratigraphically above these are schists,
metaconglomerates and amphibolites of the Vadito Group, followed by the Hondo Group. The
Hondo group has two formations, lower kilometer thick Ortega Formation and the upper Rinconada
Formation mainly composed of schists and phyllites. The Hondo Group is generally interpreted to
be a cratonic margin sedimentary sequence that can be correlated from range to range in northern
New Mexico (Bauer and Williams, 1989). This stratigraphy is cut by ~1.4 Ga and ~1.7 Ga.
granites. (Table 2.1,2.2) (Table 2.1,2.2)
63
2.3 Tres Piedras Granite
2.3.1 Occurrence
The Tres Piedras Granite is exposed around the town of Tres Piedras of eastern Rio Arriba County
(Fig. 2.1, 2.1a). It is a pink, flesh-colored moderately to well foliated medium grained granite.
Crude foliation is defined by mica rich layers. The rock -foliation strikes northwest – southeast.
The rocks in this province consist of a thick section of interbedded metasedimentary and
metavolcanic units. The major metamorphic rock types exposed at the surface are granite – gneiss,
quartzite, hornblende – chlorite schist and phyllite. Tres Piedras Granite is a group of granitic to
granodioritic bodies that intrude the metasedimentary units. The granites and the metasediments
have been metamorphosed to upper amphibolite facies and deformed during later metamorphic
events.
The exact shape and size of the granitic – gneiss body is unknown due to subsequent deformation
with Tertiary and Quaternary sediments and volcanics covering all but isolated outcrops. From the
surface expression of the outcrops it can be seen that the granite unit is elongated in shape and has
its longer axis subparallel to the regional foliation tend of the crystalline province. Large outcrops
of the Tres Piedras Granite near the town of Tres Piedras and along the Tusas River Canyon to the
west of Tres Piedras (Fig. 2.1, 2.1a). The Tres Piedras Granite near the town of Tres Piedras
consists in places of granitic inliers whose base is hidden beneath thick basalt flows related to the
Rio Grand Rifting.
2.3.2 Petrography & Chemistry
The rock is granite, ranging from very close to alkali feldspar granite to true granite. The
composition (eye estimation by volume) can be expressed as about 40% potash feldspar, 15%
plagioclase, 20% quartz; the rest 25% is made up of biotite, hornblende, muscovite, epidote,
sphene, chlorite, and rarely calcite and apatite. Some portions show weak to moderate
recrystallization texture attributed to low-grade metamorphism. The overall texture is that of
typically igneous allotriomorphic granular with most grains of felsic minerals being anhedral. Only
the biotites show a preferred orientation that define the crude foliation.
The alkali feldspar rich portions are dominated by large grains of flame and string perthite. In a few
perthite grains the exsolved lamellae of albite are oriented in two mutually perpendicular directions.
Some perthite and also some homogeneous grains of potash feldspar show thin sodic rims that in
many cases merge with the albite lamellae in perthitic grains. Graphic intergrowth of quartz are
present in some smaller grains of potash feldspar. Some portions of the rocks show strong alteration
of potash feldspar to sericite in crystallographically controlled directions. Some grains show two
such orientations. Microcline is abundant in portions, showing cross-hatched twinning. Some
microcline grains also show perthitic intergrowths. Potash feldspar grains are the largest among all
the minerals.
Plagioclase grains are more regular in shape compared to alkali feldspar and quartz grains, and
even after strong alteration to epidote and muscovite show lath shaped outline. Lamellar twinning
is characteristically present.
64
Biotite (very dark brown, and greenish brown varieties) appears to be of two different generations.
Dark brown varieties are of weakly tabular in shape and are cross cut by clean, large, perfectly
tabular muscovite grains implying the metamorphic nature of muscovite. These biotites could be
pre-metamorphic and igneous in origin. The greenish variety of biotite form composite grains with
chlorite, and green hornblende with bluish tinge has grown over the biotite. This assemblage of
chlorite – biotite – hornblende is interpreted to be metamorphic. However, most of the grains of
hornblende are irregular in shape, and some even do not show any good cleavage. Only a few
hornblende grains are subhedral with very good cleavage. These hornblende grains show brownish
pleochroism and could be of igneous origin.
Some very large muscovite grains show sieve structure defined by large sub-circular inclusions of
quartz. These seem to be older than the prismatic smaller, and inclusion free grains of muscovite.
Thus the larger grains could be of igneous origin. Quartz grains are all anhedral. Larger grains
show a tendency to form aggregates of smaller grains – but no preferred orientations are noticeable.
Lozenge-shaped sphene grains, and a few prismatic epidote grains show metamict texture defined
by dark rims around the grains. These are common around inclusions inside hornblende.
History: Alkali feldspar dominated granite with small amount of biotite, hornblende and muscovite
was metamorphosed to lower amphibolite facies.
A chemical and normative analysis of the Tres Piedras Granite by Barker (1958) is included in
Table 2.3. A comparison of the modal analysis on the same rock by Barker (1958), Bingler (1965)
and this study is included in the table too. The Tres Piedras Granite has very low content of calcium
and is weakly peralkaline (presence of normative corrundum).
2.3.3 Contact Relation
Several intrusive contacts between the Tres Piedras Granite and the surrounding rocks can be seen
in the field. Xenoliths of the country rocks are common within the Tres Piedras Granite in some
localities. Along the Tusas River Canyon, a large elongated block of muscovitic quartzite
approximately 100 meters in thickness is enclosed by granitic-gneiss. Northwest of Tusas River
Canyon a body of Tres Piedras Granite is found in which approximately one-third of the total rock
is composed of schist xenoliths (Barker, 1958). In a location southeast of the well defined contact, a
wide diffuse contact zone between granite-gneiss and muscovite quartzite is found to consist of
irregular sills which form a hybrid zone between the two units. To quote Barker (1958), “These
sills have absorbed varying amounts of muscovite quartzite and granite. The intrusion was followed
by hydrothermal reactions between wall rock and granite which resulted in hybridization of
xenolithic material and of granite along the contacts.”
2.4 Previous Work
The Proterozoic orogenic belt in the southwestern United States is a grand area to study the growth,
stabilization and reactivation of the continental lithosphere. Isotopic and geochronologic data points
towards largely juvenile crust of the orogenic belt (Bennett and DePaolo, 1987) assembled between
1.75 and 1.6 Ga. Following termination of orogenic activity ca. 1.6 Ga, the belt was amagmatic
until the intrusion of 1.48-1.32 Ga “anorogenic” granitoids in the southwestern United States
(Anderson, 1983) and 1.1-1.2 Ga mafic intrusions in the southwestern United States (Keller et al.,
65
1989). The Tres Piedras Granite along with numerous other felsic intrusions like the Embudo
Granite and the Maquinita granodiorite were intruded during the phase of juvenile crust building
between 1.75 and 1.6 Ga.
The granite exposed along the Rio Tusas and at Tusas Mountain was called the Tusas granite by
Just (1937), who also included under this name the Maquinita granodiorite and granite at Tres
Piedras. The granite exposed at Tres Piedras and along Tusas Canyon and the lower Rio Tusas
differs from that of Tusas Mountain. The granite at Tusas Mountain appears to be an atypical, fine
grained porphyritic granite. Barker (1958) redefined the granite at Tres Piedras and along the Rio
Tusas as the Tres Piedras Granite after the excellent exposure in and around that town. Hedge et al.
(1977) thought Tres Piedras Granite was probably emplaced during the same intrusive event as the
Maquinita Granodiorite as both foliated rock types bear the structural imprint of a single period of
regional metamorphism that affected their wall rocks. The age of the granodiorite emplacement was
dated by U-Pb zircon ages at 1755 m.y (Silver and Dickinson, 1987).
The Tres Piedras Granite is comparable with the Embudo Granite located in north central New
Mexico in the Picuris Range of the southern Sangre de Cristo Mountains in more than one aspect.
Both of them have intruded the Ortega Quartzite and the Vadito Group. Petrographically and
texturally they are similar. Rb- Sr whole-rock isotopic analyses indicate that the Embudo Granite
crystallized 1673 +/- 41 Ma. ago (Fullager & Shiver, 1973). This intrusive event is temporally
related to the igneous and metamorphic events and formation of regional structures involving the
tectonic emplacement of 1.65-1.7 Ga granitoids within the supracrustal rocks (Fullager & Shiver,
1973; Pedrick et al., 1998).
Maxon (1976) reported analyses of data from six Tres Piedras Granite zircon concentrations that
delineate a chord with an upper intercept on concordia at 1654 Ma. For the U-Pb analysis
performed for his study, zircons were obtained from four samples of Tres Piedras Granite type
locality (that includes the exposures around the town of Tres Piedras) and one sample from Tusas
River Canyon. In addition four zircon separates were obtained from Ortega Quartzite. The zircons
all show the usual discordance in which the age pattern is Pb207/Pb206 > Pb207/U235 > Pb206/U238. The
Ortega Quartzite samples on the concordia diagram intersects the chord (upper intercept) at 1830
Ma (Maxon, 1976).
Throughout northern New Mexico metamorphic conditions are characterized by the co-exsistence
of all three aluminosilicate polymorphs. In the Rinconada Formation, garnet-biotite thermometry
indicates temperatures of 530 +/- 30 °C at 4 kbar (Holdaway and Goodge, 1990). Currently, the
absolute timing of peak metamorphism in northern New Mexico is debated (Bauer and Williams,
1994; Grambling, 1989). The Rb-Sr whole rock and 40Ar-39Ar muscovite and hornblende mineral
ages regionally yield ages of ca. 1.4 Ga and younger, implying that much of the metamorphism
may be of this age (Mawer et al. 1989; Thompson et al. 1991; Grambling and Dallmeyer, 1993).
On the other hand, U-Pb zircon ages of ca. 1.65-1.7 Ga from plutons which cross-cut strongly
deformed supracrustal rocks are consistent with earlier fabric development. Karlstrom et al. (1994)
suggest that locally a later (1.4 Ga) 500-600 °C, 4-5 kbar metamorphism is superimposed on an
earlier (1.6 Ga) 700-800 °C, 7-9 kbar metamorphic assemblage.
66
Barker (1958) reported a pegmatitic hydrothermal event that produced numerous pegmatite
injections, including the Harding Pegmatite in north central New Mexico. Gresens (1975) related
the emplacement of the pegmatites to the regional metamorphism and associated them genetically
with metasomatically altered metarhyolite. Most pegmatite bodies are quartzofeldspathic, with the
exception of Li-rich Harding pegmatite. Montgomery (1953) assumed that the source for all
pegmatite units is the Embudo Granite, which intrudes bedded Precambrian formations. Long
(1972) dated the pegmatites of the La, Madera quadrangle to be 1425 +/- 25 Ma. His age
calculation was based on Rb-Sr isotopic data from whole rocks, metamorphic muscovite and
pegmatite “book muscovite”. Long (1972) concluded from low error on the age that the igneous
intrusion and final stages of metamorphism appear to have ceased within a time interval too short to
be resolved by the Rb-Sr method.
The staurolite and garnet separates from a sample of the Rinconada Formation from the Picuris
Mountains analyzed by Lanzirotti and Hanson (1997) yielded U-Pb ages of about 1461 ± 13 Ma.
These data are consistent with metamorphism at 1450 Ma in northern New Mexico which results in
porphyroblast growth.
Table 2.1. The regional geologic history of the north-central New Mexico is summarized in the
following table.
AGE in Ma.
GEOLOGIC EVENT
~1425
Pegmatitic event following the regional metamorphism. (Long, 1972; Gresens,
1975)
~1460
Regional metamorphism following the felsic intrusions. Peak metamorphic
temperature 400-500 degrees C at 4-5 Kb pressure.(Lanzirotti and Hanson, 1997)
1480-1400
Anorogenic granitoid intrusion (Anderson, 1983, Bowring and Karlstrom, 1990)
1600
Regional metamorphism following the granite intrusion. Peak metamorphic
temperature 700-800 degrees C at 7-9 Kb pressure.(Grambling and Dallmeyer,
1993)
Granites and granodiorites like Maquinita granodiorite, Embudo “type” granite
intruding the mafic complex and/or the metasedimentary sequence (Fullager and
Shiver, 1973; Gresens, 1975). Tres Piedras Granite intrusion (Maxon, 1976)
Deposition of the Hondo Group. It is thought to be a cratonic margin sedimentary
sequence. (Bingler, 1974)
1650-1750
<1700
1700
Deposition of the Vadito Group.(Bingler, 1974)
1720-1765
Moppin, Pecos, Gold Hill and similar mafic complex geochemically similar to
modern tholeiitic and/or calc alkaline rock found in arc or back arc assemblage
(Gabelman, 1988)
67
Table 2.2 Stratigraphic nomenclature and lithologic description of supracrustal Proterozoic rocks
of New Mexico. Modified after Bauer and Williams, 1989.
Nomenclature
Description of Lithology
Age (Ma)
HONDO GROUP
Piedra Lumbre Formation
Pilar Formation
Phyllites
Rinconada Formation
Schists and quartzite.
Ortega Formation
Thick orthoquartzite.
VADITO GROUP
The unit is divided into a lower member dominated
by conglomerate with minor interlayered micaceous
quartzite and an upper member dominated by
micaceous and feldspathic quartzite.
Burned Mt. Formation
Massive quartz-feldspar rock characterized by
distinctive quartz and feldspar eyes in a fine grained
laminated matrix.
Marquenas Formation
Consists of approximately equal amount of
metaconglomerate and quartzite.
Glenwoody Formation
Approximately 1000 ft of feldspathic schist
containing quartz and feldspar megacrysts.
MAFIC METAVOLCANIC SEQUENCE
1700 Ma.
Pecos Complex
1720 Ma.
U-Pb zircon dating of
metavolcanic
rocks
(Bowring and Condie,
1982, Robertson and
Condie, 1989).
Big Rock Formation
Gold Hill Complex
Moppin Complex
Consists of 1. metavolcanic rocks with minor
interbedded volcaniclastic rocks and iron formation;
2. metasedimentary rocks including feldspathic
sandstone, shale, volcaniclastic graywacke, carbonate
and distal iron formation; 3. Intrusive felsic and
mafic rocks of a subvolcanic complex; 4. Younger
granitic rocks.
A mafic/felsic layered gneiss succession consisting
of complexly interlayered feldspathic gneiss, biotitehornblende gneiss, hornblende gneiss and
amphibolite. The sequence is interpreted as
metamorphosed volcaniclastic rocks interlayered
with flows, tuffs and sedimentary rocks.
Mafic metavolcanic sequence with chlorite schist
and/or amphibolite with lesser components of
feldspathic schist and gneiss, muscovite (+/-biotite)
schist, metaconglomerate and BIF. Locally exposed
primary volcanic and sedimentary structures include
pillows, graded beds and relict phenocrysts. The
Moppin complex is associated with (intruded by?)
several granodiorite bodies collectively called the
Maquinita Granodiorite.
68
Metavolcanic rocks from
several localities within
the Vadito Group have
yielded U-Pb zircon
isotopic
ages
of
approximately 1700 Ma.
(Silver,
1984
as
referenced in Bauer and
Williams, 1989)
1765 Ma.
U-Pb zircon data from a
felsic metatuff near Gold
Hill (Bowring, 1986)
Older than 1755 Ma.
U-Pb age of zircons from
Maquinita Granodiorite is
(Silver and Dickinson
1987)
2.5 Results
2.5.1 Rb-Sr Whole Rock Analysis
In the Rb-Sr whole rock analysis of the of the Tres Piedras Granite (Table 2.4) , the following
samples were analyzed: eleven samples from the Tres Piedras Granite type locality (near the town
of Tres Piedras) and five samples from the Tres Piedras Granite exposed along the Tusas River
Canyon. The ages determined by Isoplot 3 (Ludwig, 2001) and plotted in the cubic least square
isochron diagram were as follows: Tres Piedras Granite Type Locality – 1490 Ma ± 20 Ma with an
initial Sr87/Sr86 ratio of 0.7182 ± 0.0006; Tres Piedras Granite at Tusas River – 1497 Ma ± 42 Ma
with an initial Sr87/Sr86 ratio of 0.7147 ± .0012 (Figure 2.2a, b, c). The data indicates that the RbSr system of the Tres Piedras Granite is not much disturbed and the data delineate good isochrons
(MSWD <3) for the both the bodies studied.
2.5.2 Rb-Sr Mineral isochron
Analyses were made on total rock samples and mineral separates including sphene, feldspar and
biotite from Tres Piedras Granite type locality sample and Tusas River Canyon sample. Using the
Rb-Sr isochron diagram, the measured ratios 87Sr/86Sr and 87Rb/86Sr are plotted for the two
locations. Analytical error is less than the size of the symbol. If all of the minerals have reequilibrated at the time of metamorphism and were subsequently closed they would all lie on a
common straight line.
The data for the Tres Piedras Granite type locality sample is shown in Figure 2.3, Table 2.5 . The
biotite-whole rock-sphene isochron corresponds to an age of 1472±40 Ma with an initial 87Sr/86Sr
of 0.724±0.037 with feldspar lying above the isochron (Figure 2.3a). The deviation of the feldspar
from the biotite-whole rock-sphene from the Tres Piedras type locality sample may represent
incomplete re-equilibration or possibly due to later contamination by radiogenic strontium and/or
rubidium loss.
The Rb-Sr mineral isochron obtained from the Tusas River Canyon locality yielded an isochron of
1509±40 Ma with initial 87Sr/86Sr =0.731±0.031 (Figure 2.3b). In this case also the feldspar did
not fall on the biotite-whole rock-sphene isochron and was lying above the isochron.
2.5.3 U – Pb zircon ages by LA-MS-ICPMS.
Zircon were separated from the Tres Piedras Granite type locality and Tusas River Canyon samples
and analyzed with LA-MS-ICPMS facility at the University of Arizona (Table 2.6). When plotted
on the concordia diagram 20 zircon analyses from the Tres Piedras type locality samples yielded a
chord with the upper intercept at 1732.7±4.2 Ma (Figure 2.4a). 29 zircon measurement from the
Tusas River Canyon locality delineates a chord in the concordia diagram with the upper intercept at
1729±20 Ma and lower intercept at 90±20 Ma (Figure 2.4b). The zircon populations of the Tusas
River Canyon are affected by lead loss more than the Tres Piedras type locality samples. Thus the
age of crystallization of the Tres Piedras Granite is ~1730 Ma as evidenced by the LA-MS-ICPMS
analysis of the zircons separated from the granite samples from two different localities.
69
2.6 Discussion
The U-Pb zircon ages of the Tres Piedras Granite and Tusas River Canyon are significantly
different from the Rb-Sr whole rock ages in both the locations-Tres Piedras Granite type locality
and Tusas River Canyon. The possible explanations for this difference can be:
A) The Rb-Sr whole rock age (~1470 Ma) is the crystallization age of the Tres Piedras Granite and
the isotope dilution U-Pb age (~1654 Ma) of zircons (Maxon, 1976) are due to inheritance of
xenocrysts.
If the Rb-Sr whole-rock age is the true crystallization age of the Tres Piedras Granite, then the 1654
Ma U-Pb age of the zircons might be biased due to incorporation of older zircon xenocrysts. This
could occur if 1500 Ma granite assimilated wall rock during emplacement. In that case the 1654 Ma
age may represent a “mixing” age between younger 1500 Ma zircon and unreset or partly reset
older zircons. The rocks which are in intrusive contacts with the Tres Piedras Granite are
predominately quartz-muscovite schists which are believed to be metasomatized rhyolites (Gresens,
1971), and a pluton of calc-alkaline granodiorite (Maquinita Granodiorite). Barker, 1974
determined that the ages of these bodies fall into the 1750-1800 Ma age range as determined by UPb zircon ages. Tres Piedras Granite that intruded the 1890 Ma Ortega Formation (Maxon, 1976)
could also assimilate small amount of the quartzite. Therefore these rocks provide a source for the
zircons that could be later incorporated into the possibly 1500 Ma Tres Piedras Granite. This
problem can be easily resolved by looking at the single zircon ages.
Among the 20 zircons analyzed by laser ablation ICPMS in the Tres Piedras Granite sample more
than 18 grains are between 97 to 100 % concordant (difference between 206Pb/238U age and
206
Pb/207Pb age). The presence of a high percent of concordant zircon between 1720 and 1730 Ma
excludes the possibility of incorporation of xenocrystic zircons. In that case some type of
disturbance must have occurred in the Rb-Sr isotopic systems of the Tres Piedras Granite to cause
the isochron ages to differ from the zircon ages. This might give rise to two possibilities for this
disturbance in the Rb-Sr whole rock system of the Tres Piedras Granite.
B) Migration of the isotopes (open system) and lowering of the age by this migration.
In an open system, with the lowering of age, there is a need to maintain isochron linearity and the
process must involve the migration of either both Rb and Sr or loss or gain of equal amount of Sr or
Rb respectively from each sample. The isochron obtained for the whole rock samples show a good
linearity of data points. This low spread in the data (MSWD < 3.0) suggests that the change in the
87
Sr/86Sr or the 87Rb/86Sr ratios of each sample due to isotope exchange must be closely
proportional to the particular Rb/Sr ratio of the sample. This would require either the addition of a
proportional amount of Rb and the removal of a proportional amount of Sr from all samples or
both. It is doubtful whether trace element metasomatism would proceed in such an orderly manner
and the close alignment of the data points makes implausible demands upon a transport mechanism.
C) The internal redistribution or isotopic homogenization at approximately 1500 Ma.
70
Tres Piedras Granite crystallized at ~ 1700 Ma but was later affected by regional metamorphism at
~1500 Ma that redistributed and homogenized the Sr isotope. The high initial 87Sr/86Sr ratio (0.7182
and 0.7147) and good isochron linearity indicate that this homogenization at approximately 1500
Ma was on a scale larger than the size of the whole-rock samples and was at least the size of the
domain over which the samples were collected (approx. 50 sq. mt. of collecting area at each
exposure). The best fit line yields an age that seem to correspond well with Lanzirotti and Hanson’s
(1997) staurolite and monazite age of amphibolite grade metamorphism between at 1480 Ma. from
the Picuris range. If each collecting location represents a single domain in which the whole-rock
isochron has undergone resetting due to metamorphism, the data then appear to indicate that limits
can be set on the extent of isotopic homogenization for the Tres Piedras Granite.
D) Statement of preferred interpretation.
If the Tres Piedras Granite has undergone isotopic homogenization on the outcrop scale, it should
show similar behavior on the mineral scale too. Both the Tres Piedras Granite sample and the Tusas
River Canyon sample define a biotite-whole rock-sphene isochron of 1472±40 Ma and 1509±40
Ma respectively. The initial 87Sr/86Sr of the two locations are similar within error (0.724±0.027 and
0.731±0.031). Thus isotopic homogenization of between 1470 and 1500 Ma is evidenced in the RbSr whole rock as well as the biotite-whole rock- sphene isochron.
Feldspar analyzed in both the samples fall off the isochron and lies above it in both cases. The
deviation of the feldspar grains from the biotite-whole rock-sphene isochron is possibly due to
incomplete equilibration of the feldspar or later contamination.
Thus if the feldspar remained an open system in the Tres Piedras type locality with respect to
rubidium and/or strontium it might have developed a concentration gradient with respect to
rubidium and/or strontium. In order to check any inhomogenity in the feldspar grains (separated
from Tres Piedras type locality sample), the Rb and Sr concentrations were measured across single
feldspar grains by laser ablation in Finnigen element ICP-MS. Several grains were measured and
the Rb/K, Sr/K and Rb/Sr ratio of one traverse of a single grain are shown in Figure 2.5. None of
the grains showed any concentration gradient from core to rim. The local variability of the ratios
can be attributed to a small degree of sericitization of feldspars. In the absence of any concentration
gradient of Rb and Sr it was necessary to evaluate whether or not the feldspars were homogenized
with respect to Sr isotope.
Approximately 100 mg of feldspar grains with well developed crystal faces were hand picked under
a binocular microscope. They were partially dissolved in cold distilled 3:1 HF-HNO3 acid for 15
minutes which dissolved 20 % (by weight) of the grains. This dissolved the rims of the feldspar
grains and the remaining partially dissolved grains were similarly treated with distilled 3:1 HFHNO3. The core and rim solutions of the feldspars were passed through the cation column to
separate Sr that was measured for the 87Sr/86Sr ratio in the TIMS.
The 87Sr/86Sr ratio obtained for the rim is 0.88429 ± 0.00011 and that for the cores is 0.85071 ±
0.00011 indicating that with respect to the Sr isotopic ratio the feldspar grains are inhomogeneous,
possibly causing the feldspar to deviate from the 1472 Ma mineral isochron of the Tres Piedras type
locality sample. Assuming that the feldspars were homogenized at the 1500 Ma metamorphic
71
event, the Δ87Sr/86Sr of the rim and the core observed at present can be caused by a difference of
0.56 in the Rb/Sr ratio of the rim and the core at 1500 Ma.
Discordance in the Rb-Sr whole rock age and U-Pb zircon ages of granitic plutons are common in
the literature, but the cause of this difference is rarely complete isotopic homogenization of Sr
isotope during a later metamorphic or metasomatic event. In most cases the discordant Rb-Sr whole
rock age is explained by resetting, which rarely corresponds to the regional metamorphic age. In
other cases the U-Pb isotope dilution zircon ages are older than the true crystallization age due to
incorporation of older xenocrystic zircons. Examples include the Crossnore Complex of the
southern Appalachians (e.g. Davis et al., 1962; Rankin et al., 1969; Odom and Fullager, 1984),
World Beater Complex of California (Lanphere et al., 1963), Mount Isa of Australia (Page,
1987)etc.
Rb-Sr data on the Neoproterozoic alkalic to peralkaline rift related granites of the Crossnore
plutonic suite yielded and age between 681-706 Ma (Odom and Fullager, 1984) whereas the zircon
ages ranged between 690 and 820 Ma (e.g. Davis et al., 1962; Rankin et al., 1969; Odom and
Fullager, 1984). It was established that there exists at least two zircon populations of different ages
(Odom and Fullager, 1984) and more recent studies (Su et al., 1994) determined the age of the
Crossnore plutonic complex to be 741 Ma, with inherited pre-Grenville source components of 1424
Ma that corresponds well with the whole rock Nd model ages of 1300-1500 Ma. In the case of the
Tres Piedras Granite the laser ablation study of single zircon clarified the absence of any inherited
zircon xenocrysts that might have caused the difference in age between the Rb-Sr whole rock and
U-Pb (isotope dilution) age of the zircons.
Page (1978) conducted a detailed study of the Rb-Sr whole-rock and mineral and U-Pb zircons of
the Proterozoic silicic volcanic rocks and granitic intrusions from near Mt. Isa, northwest
Queensland, Australia. U-Pb zircon ages within the basement igneous succession show that the
oldest recognized crustal development was the out pouring of acid volcanics at 1865 Ma which
were intruded by coeval granites and granodiorites (Kalkadoon Granite) whose U-Pb age is ~1862
Ma. All of these rocks are altered in various degrees by low grade metamorphic events that can be
related to the emplacement of a syntectonic granite batholith (Wonga Granite) at 1670 Ma ago. The
rocks that significantly predate this earliest recognized metamorphism have had their primary RbSr whole rock systematics profoundly disturbed as evidenced by 10 to 15 % lowering of most RbSr isochron ages and many of the lowered ages (some of which are in conflict with unequivocal
geological relationships) are within the 1600-1700 Ma interval. Page (1978) argued that the
processes that modified the Rb-Sr whole rock ages are complex and may have been related to the
first major greenschist metamorphism (1670-1625 Ma) and a younger greenschist facies event
(1620-1490 Ma) superimposed further isotopic re-distribution for Rb-Sr isotopes. The Kalkadoon
Granite gave different initial 87Sr/86Sr ratios for different sets of mineral isochrons thus producing
parallel isochrons for samples from different locations of the same pluton. Thus, Page attributed
this to source characteristics rather than a secondary incomplete homogenization feature. He
explained the parallel isochrons with different initial ratios produced by different pulses of
magmatism in the same granite body.
The initial 87Sr/86Sr ratios obtained from the whole-rock analysis of Tres Piedras Granite and Tusas
River Canyon range from 0.7147 to 0.7182 (though they are indistinguishable within error margin).
72
The two 87Sr/86Sr ratios observed in the Tres Piedras Granite can be generated in 250 Ma with a
87
Rb/86Sr ratio of 3 and 4 respectively with a starting 87Sr/86Sr ratio of 0.704 of the primary
magma. If we assume a fixed elemental ratio of 87Rb/86Sr =2.8 (average of all the analyzed whole
rocks) the initial 87Sr/86Sr ratio required are 0.705 and 0.708 respectively.
Partial isotopic homogenization of the minerals during a metamorphic event is common also. In the
World Beater Complex of Paramint Range in California the Precambrian and Paleozoic rocks have
been metamorphosed in late Mesozoic time. K-Ar ages on biotite gives ages ranging from 103-130
Ma (Lanphere et al., 1963). Rb-Sr isotopic studies of all the constituent minerals (apatite,
plagioclase, K-feldspar, muscovite and biotite) and their associated total rock yielded biotite-total
rock isochrons indicating ages ranging from 64 to 156 Ma (Lanphere et al., 1963). In each case
apatite and muscovite was lying below and above the isochron and the observation was explained
by incomplete isotopic homogenization of these two mineral phases.
Discrepancy in age Rb-Sr whole rock and U-Pb zircon age is also recorded in the granites and
rhyolites of St Francosis Mountain of southeastern Missouri. The U-Pb zircon ages are in the range
of ~1530 Ma where as the Rb-Sr whole rock ages are between 1273 Ma to 1408 Ma (Bickford and
Odom, 1969; Mose and Bickford, 1972). Initial 87Sr/86Sr of all the rocks are characterized by large
errors caused in part by scatter of analytical points (effect more pronounced particularly for the
volcanic rocks) but mostly by the high 87Rb/86Sr ratios that are so characteristics of these rocks.
These uncertainties in the initial ratios do not appreciably affect ages computed from the isochron,
but they make the initial 87Sr/86Sr ratios virtually useless for petrogenetic interpretations. Rb-Sr
ages of the volcanic rocks show a significant positive correlation with Sr concentrations (Mose and
Bickford, 1972) suggesting that Sr loss is responsible for lowering the Rb-Sr age determination.
Thus it was concluded by Bickford and Mose (1975) that the igneous rocks of the St Francosis
Mountain were formed at ~ 1500 Ma and some subsequent event disturbed the Rb-Sr system. As
the rocks were not affected by regional metamorphism the authors postulated that Sr loss could
have occurred during widespread hydrothermal alteration of the rocks, an event suggested by the
generally turbid feldspar, alteration of mafic minerals and occurrence of epidote. A hydrothermal
event is also suggested by iron mineralization in the area as typically hematite and magnetite
replace the volcanic rocks.
Though there is evidence of hydrothermal event (pegmatite injections) in the north central New
Mexico at ~1425 Ma but the Tres Piedras Granite do not show any sign of hydrothermal alteration
either in hand specimen or under the microscope.
Diversity of age clusters obtained from regional metamorphic areas are very often interpreted as the
result of isotopic homogenization that causes the "resetting" and "rejuvenation" of age values.
Largescale isotopic homogenization is rarely attained; and when it does, it requires extensive
reaction and chemical exchange with a pervasive fluid representing one end member of fluid-rock
interaction. During metamorphism fluids result due to dehydration reactions and are quickly
removed from the system pervasively in case of permeable rocks and through cracks and fractures
in impermeable or less permeable rocks. It is more common that fluid – rock interaction will be in
local scale and patchy (Cartwright and Valley, 1992; Rumble et al., 1983; Rye et al., 1976;
Sheppard and Schwarcz, 1970; Valley and O'Neil, 1984) rather than to be in regional scale.
73
The Tres Piedras Granite has been affected by regional metamorphism and there is evidence of
alkali metasomatism during the intrusion of pegmatite (Bingler, 1974). Regional metamorphism is
the most widespread of any type of metamorphism occurring over broad expanses of deeper levels
of the crust that involves reconstitutive changes in fabric and composition of the rock body.
Reconstitutive metamorphic changes occur in the solid state, in the presence of usually minor
amounts of aqueous or carbonic fluids. If the homogenization process of the Tres Piedras Granite
represents such a solid state reconstitution it will be interesting to study the O-isotope partition as
that might suggest the extent to which Si, Al, O participated in the homogenization and whether the
process should be regarded as an exchange process or a recrystallization process.
2.7 Conclusion
A fundamental problem in geochronology in the interpretation of the discordant mineral and total
rock ages which are commonly obtained in areas which have had a polymetamorphic histrory.
However the age discordance itself provides a valuable tool in the study of the petrogenesis and
evolution of a polymetamorphic terrane.
There can be three possibilities that can explain the discordance of the Rb-Sr whole-rock age and
the U-Pb zircon age: A) the U-Pb zircon ages is biased due to incorporation of older zircons B) the
Tres Piedras Granite was an open system and Rb-Sr whole rock age is lowered by the migration of
the element(s) C) internal redistribution or isotopic homogenization of the Sr isotopes at 1500 Ma.
U-Pb analysis of single zircons by Laser ablation MC-ICPMS and presence of a large number of
concordant zircons between 1720-1730 Ma excluded the possibility of presence of xenocrystic
zircons. If the U-Pb zircon ages of ~1720 Ma indicate the time of crystallization, then some type of
disturbance must have occurred in the Rb-Sr isotopic systems of the Tres Piedras Granite to cause
the type locality and Tusas River Canyon isochron ages to differ from the zircon age. Open system
behavior is not favored as that would require equal loss or gain of Sr and Rb respectively from all
the samples. The good linearity of the isochron with minimum spread of the data along with the
high initial 87Sr/86Sr ratio favors the possibility for this disturbance is internal redistribution or
isotopic homogenization at 1500 Ma.
The high initial ratio and good isochron fit of the whole rock samples of Tres Piedras Granite from
two locations indicate that this homogenization at approximately 1500 Ma. was on a scale larger
than the size of the whole rock samples and was atleast the size of the domain over which the
samples were collected (approx. 100 sq. mt. of collecting area at each exposure). The mineral
isochron in both the locations corresponds well with the whole-rock Rb-Sr age. The strontium
isotopic composition of sphene and biotite was almost completely homogenized. The degree of
strontium isotopic homogenization is less complete for feldspar. The 1500 Ma. isotopic
homogenization age seems to correspond well with Lanzirotti and Hanson (1997) staurolite and
monazite age of amphibolite grade metamorphism at 1480 Ma. from Picuris range. If each location
represents a single domain in which the whole rock isochron has undergone resetting due to
metamorphism, the data then appears to indicate that limits can be set on the extent of isotopic
homogenization for the Tres Piedras Granite.
74
Table 2.3. Chemical analysis, norm and modes of the Tres Piedras Granite.
Chemical Analysis
(Barker, 1958)
Norm
(Barker, 1958)
Tres Piedras Granite
(n=16)
Tres Piedras Granite
SiO2
TiO2
Al2O3
Fe2O3
FeO
MnO
MgO
CaO
Na2O
K2O
P2O5
H2O
CO2
TOTAL
77.25
0.14
11.56
0.55
0.92
0.03
0.11
0.27
2.65
5.59
0.06
0.53
0
99.66
Quartz
Orthoclase
Albite
Anorthite
Magnetite
Ilmenite
Enstatite
Ferrosilite
Corundum
TOTAL
39.5
33.5
22.5
1.1
0.7
0.3
0.3
1.1
0.6
98.6
Modes
Quartz
Microcline
Albite
Muscovite
Biotite
Total
75
Barker, 1958 Bingler, 1965
This Study
(n=6)
This Study
(n=5)
Tres Piedras
Granite
Tres Piedras
Granite
Tres Piedras
Granite
43
38
13
3
3
100
47
16
23
10
4
100
44
31
24
1
100
Tusas
River
Canyon
45
43
12
1
101
Table 2.4 Rb-Sr isotopic data obtained by mass spectrometric analysis of whole rock
samples of the Tres Piedras Granite.
SAMPLE
Rb, ppm
Sr, ppm
TPGT 1
TPGT 2
TPGT 3
TPGT 4
TPGT 5
TPGT 6
TPGT 7
TPGT 8
TPGT 9
LT-02-05
LT-02-4B
TPGT 11
TPGT 12
TPGT 13
TPGT 14
TPGT 15
139.6
55.7
68.3
88.6
145.5
64.1
63.9
180.9
71.1
140.5
20.1
124.6
109.2
74.5
128.9
155.4
177.2
135.7
183.1
152.1
165.3
158.6
158.4
85.8
150.0
60.17
224.3
156.5
158.3
189.6
130.3
162.9
Rb87/Sr86
2.250 ± 0.0007
1.169 ± 0.0010
1.062 ± 0.0007
1.661 ± 0.0005
2.515 ± 0.0030
1.150 ± 0.0009
1.149 ± 0.0014
6.090 ± 0.0012
1.350 ± 0.0014
6.538 ± 0.0005
0.301 ± 0.0013
2.272 ± 0.0025
1.958 ± 0.0012
1.119 ± 0.0007
2.827 ± 0.0005
2.725 ± 0.0005
Sr87/Sr86
0.7669
0.7432
0.7410
0.7526
0.7706
0.7406
0.7427
0.8455
0.7473
0.8593
0.7228
0.7623
0.7563
0.7381
0.7740
0.7722
Sr/86Sr measured ratio normalized to 86Sr/88Sr = 0.1194.
Precision of replicate whole rock determinations is 0.001% for 87Sr/86Sr.
87
Rb – Sr Whole Rock Ages
Location
Type Locality
Tusas River Canyon
Age
1490 +/- 20 m.y
1497 +/- 42 m.y
(Sr87/Sr86)0
0.7182 +/- 0.0006
0.7147 +/- 0.0013
76
± 0.0008
± 0.0010
± 0.0007
± 0.0008
± 0.0006
± 0.0006
± 0.0009
± 0.0003
± 0.0004
± 0.0004
± 0.0009
± 0.0008
± 0.0006
± 0.0004
± 0.0003
± 0.0005
LOCATION
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Type Locality
Tusas River Canyon
Tusas River Canyon
Tusas River Canyon
Tusas River Canyon
Tusas River Canyon
Table 2.5 Rb-Sr isotopic data obtained by mass spectrometric analysis
of mineral phases of the Tres Piedras Granite.
Rb87/Sr86
Sr87/Sr86
Whole Rock
Sphene
Feldspar
Biotite
6.538 ± 0.0005
0.0496 ± 0.0015
1.1298 ± 0.0010
43.437 ± 0.0019
0.85934 ± 0.0004
0.72753 ± 0.00011
0.88281 ± 0.00021
1.64287 ± 0.00036
Type Locality
LT-02-05
Whole Rock
Sphene
Feldspar
Biotite
1.616 ± 0.0003
0.0869 ± 0.0019
0.6662 ± 0.0020
45.642 ± 0.0012
0.78519 ± 0.00018
0.73495 ± 0.00011
0.83400 ± 0.00021
1.74287 ± 0.00035
Tusas River Canyon
LT-02-2
SAMPLE
*87Sr/86Sr normalized assuming 86Sr/88Sr = 0.1194
Rb – Sr Mineral Isochron Ages
Location
Type Locality
Tusas River Canyon
Age
1472±40 Ma
1509±40 Ma
(Sr87/Sr86)0
0.724±0.037
0.731±0.031
77
LOCATION
Table 2.6 Tres Piedras Granite Type Locality Zircon analysis by LA-MC-ICPMS
CONCENTRATIONS AND RATIOS
U (ppm) 206Pb U/Th 207Pb* ±(%) 206Pb* ±(%) Error
204
235
238
Pb
U
U
corr
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
6534
970
469
1866
1173
2255
1709
824
3050
4845
393
1270
563
1166
1595
1131
1185
665
1357
221
36729
16625
6226
18184
13776
9443
8253
10174
48107
47685
14182
27442
13121
58470
54324
23449
12221
23459
25612
5957
6.2
4.8
4.2
6.0
6.1
4.3
4.8
4.7
4.4
5.0
2.7
5.3
5.1
4.3
6.1
4.0
5.9
4.6
5.3
1.6
4.4560
4.4470
4.4389
4.4732
4.4396
4.4486
4.4822
4.4605
4.5093
4.5139
4.5318
4.4704
4.4517
4.6671
4.4458
4.6495
4.4481
4.4458
4.4625
4.4105
1.90
2.24
2.15
1.54
1.78
1.81
2.37
2.54
2.25
1.75
3.57
1.98
1.72
1.57
1.45
2.25
1.42
3.66
2.43
1.43
0.3045
0.3041
0.3051
0.3034
0.3058
0.3061
0.3054
0.3053
0.3080
0.3080
0.3101
0.3075
0.3049
0.3185
0.3012
0.3172
0.3049
0.3017
0.3034
0.3027
1.00
2.00
1.79
1.17
1.46
1.49
2.14
2.33
2.01
1.43
3.42
1.71
1.39
1.21
1.05
2.02
1.00
3.52
2.21
1.00
0.53
0.89
0.83
0.76
0.82
0.82
0.90
0.92
0.90
0.82
0.96
0.86
0.81
0.77
0.72
0.90
0.71
0.96
0.91
0.70
206
Pb
238
U
1713.7
1711.9
1716.9
1708.4
1720.4
1721.5
1718.4
1717.5
1731.1
1731.3
1741.5
1728.7
1715.8
1782.6
1697.7
1776.4
1716.0
1699.7
1708.3
1705.0
APPARENT AGES (Ma)
±(Ma) 207Pb ±(Ma) 206Pb ±(Ma)
235
207
U
Pb
15.0
30.1
27.0
17.6
22.0
22.5
32.3
35.1
30.5
21.7
52.2
25.9
20.9
18.8
15.7
31.4
15.1
52.6
33.2
15.0
1722.8
1721.1
1719.6
1726.0
1719.8
1721.5
1727.7
1723.7
1732.7
1733.5
1736.8
1725.5
1722.0
1761.4
1720.9
1758.2
1721.3
1720.9
1724.0
1714.3
15.7
18.5
17.8
12.8
14.8
15.0
19.7
21.1
18.7
14.5
29.7
16.4
14.3
13.1
12.0
18.8
11.8
30.4
20.1
11.9
1734
1732
1723
1747
1719
1721
1739
1731
1735
1736
1731
1722
1730
1736
1749
1737
1728
1747
1743
1726
30
18
22
18
19
19
19
19
18
18
19
18
19
18
18
18
18
18
18
19
U concentration has an uncertainty of ~15%.
Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma.
Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207Pb/204Pb.
78
Table 2.6 continued.Tusas River Canyon Zircon analysis by LA-MC-ICPMS
U (ppm)
1
2
3
4
5
6
7
8
9
10
11
12
13
14
15
16
17
18
19
20
21
22
23
24
25
26
27
28
29
200
2489
843
969
2110
1154
1286
294
1272
442
1836
209
201
361
1159
2107
377
180
755
1496
468
571
425
1251
1321
2743
647
933
446
CONCENTRATIONS AND RATIOS
Pb U/Th 207Pb* ±(%) 206Pb* ±(%)
204
235
238
Pb
U
U
Error
corr
5893 2.4 4.3267 2.54 0.3004 1.00
1049 16.6 1.1977 2.96 0.0937 2.68
2663 3.6 4.4154 8.63 0.3005 8.35
1917 3.1 2.4435 4.15 0.1687 3.88
1290 2.3 1.3191 2.16 0.0940 1.89
2078 3.0 2.3748 2.62 0.1628 2.41
1411 4.8 1.6717 3.94 0.1238 3.74
22711 3.7 4.4075 1.64 0.3005 1.30
884
3.2 1.7329 1.93 0.1191 1.41
6619 3.8 3.9825 3.14 0.2690 2.92
968
2.4 1.2998 2.84 0.0949 2.54
17484 3.8 4.2284 3.24 0.2956 2.28
22725 3.0 4.3606 2.59 0.2985 2.07
8388 3.1 4.4141 1.77 0.3015 1.00
3788 4.2 4.2985 3.87 0.2984 3.73
1864 4.0 1.3994 3.56 0.0951 2.74
33317 3.5 4.3879 1.59 0.2966 1.04
22135 3.4 4.2764 2.35 0.2997 1.26
4301 2.6 3.2171 2.17 0.2193 1.91
914
3.6 1.6348 3.28 0.1151 2.73
25241 4.0 4.4092 2.15 0.2985 1.90
8897 4.2 4.4012 8.29 0.3017 8.22
6057 3.9 4.4613 3.32 0.3062 3.16
1254 2.7 2.2571 2.36 0.1524 1.98
2441 3.0 2.3126 2.52 0.1626 2.27
902
4.0 1.4066 2.51 0.1067 1.43
4189 3.3 4.3407 3.61 0.3004 3.47
1754 3.3 2.2604 8.82 0.1576 8.65
9223 3.7 3.8949 2.04 0.2685 1.76
0.39
0.91
0.97
0.93
0.88
0.92
0.95
0.79
0.73
0.93
0.90
0.70
0.80
0.56
0.96
0.77
0.65
0.54
0.88
0.83
0.88
0.99
0.95
0.84
0.90
0.57
0.96
0.98
0.86
206
206
Pb
238
U
1693.7
577.5
1694.2
1004.9
579.2
972.6
752.7
1694.2
725.8
1536.2
584.7
1669.7
1684.2
1698.8
1683.8
586.0
1674.8
1690.0
1278.3
702.5
1684.3
1699.9
1722.2
914.7
971.7
653.9
1693.5
943.4
1533.4
APPARENT AGES (Ma)
±(Ma) 207Pb ±(Ma) 206Pb ±(Ma)
235
207
U
Pb
14.9
14.8
124.4
36.1
10.5
21.8
26.6
19.4
9.7
39.9
14.2
33.5
30.7
14.9
55.3
15.4
15.3
18.7
22.2
18.2
28.2
122.8
47.8
16.9
20.5
8.9
51.7
75.9
24.0
1698.5
799.6
1715.2
1255.5
854.2
1235.1
997.8
1713.8
1020.8
1630.6
845.6
1679.6
1704.9
1715.0
1693.1
888.7
1710.1
1688.8
1461.3
983.7
1714.1
1712.6
1723.8
1199.0
1216.2
891.7
1701.1
1200.0
1612.6
21.0
16.4
71.6
29.9
12.5
18.7
25.0
13.6
12.4
25.5
16.3
26.6
21.4
14.7
31.9
21.1
13.1
19.4
16.8
20.7
17.8
68.7
27.5
16.6
17.8
14.9
29.8
62.1
16.4
1704
1482
1741
1715
1657
1728
1584
1738
1722
1755
1611
1692
1730
1735
1705
1743
1754
1687
1738
1679
1751
1728
1726
1755
1681
1539
1711
1697
1718
43
24
40
27
19
19
23
18
24
21
24
43
29
27
19
42
22
37
19
33
18
19
18
23
20
39
19
31
19
U concentration has an uncertainty of ~15%.
Isotope ratios are corrected for Pb/U fractionation by comparison with standard zircon with an age of 564+/-4 Ma.
Initial Pb composition interpreted from Stacey and Kramers (1975), with uncertainties of 1.0 for 206Pb/ 204Pb and 0.3 for 207Pb/204Pb.
79
Figure 2.1. Map of Proterozoic
basement
uplifts
and
associated major faults in
north-central New Mexico.
After Bauer and Williams,
1989.
Figure 2.1a. Map of Tres
Piedras Granite sampling
locations (Modified after
Maxon, 1976)
80
Whole Rock Isochron of Tres Piedras Granite from the Type Locality
0.86
0.84
TPGT 8
0.82
LT-02-05
87Sr/86Sr
0.80
0.78
TPGT 5
TPGT 1
0.76
Age = 1490±20 Ma
Initial 87Sr/86Sr =0.7182±0.0006
MSWD = 2.2
0.74
LT-02-4B
0.72
0.70
0
2
87Rb/86Sr
4
6
0.76
TPGT 4
TPGT 9
TPGT 7
TPGT 3
TPGT 2
TPGT 6
87Sr/86Sr
0.74
0.72
Age = 1490±20 Ma
Initial 87Sr/86Sr =0.7182±0.0006
MSWD=2.2
0.70
0.0
0.4
0.8
1.2
1.6
2.0
2.4
2.8
87Rb/86Sr
Figure 2.2. Whole rock Rb-Sr isochron of Tres Piedras Granite type locality samples.
Isochron plotted using Isoplot 3.0 (Ludwig, 2001)
81
Whole Rock Isochron of Tres Piedras Granite from Tusas River Canyon
0.82
Age = 1497±42 Ma
Initial 87Sr/86Sr =0.7147±0.0012
MSWD = 0.49
87Sr/86Sr
0.80
0.78
TPGT 14
0.76
TPGT 15
TPGT 11
TPGT 12
0.74
TPGT 13
0.72
0.70
0.0
0.4
0.8
1.2
1.6
87Rb/86Sr
2.0
2.4
2.8
Figure 2.2c. Whole rock Rb-Sr isochron of Tusas River Canyon samples.
Isochron plotted using Isoplot 3.0 (Ludwig, 2001)
82
1.9
1.7
Biotite
87
Sr/86Sr
1.5
1.3
1.1
Feldspar
0.9
Tres Piedras Type Locality
Age = 1472±40 Ma
Initial 87Sr/86Sr =0.724±0.037
MSWD = 9
WR
0.7
Sphene
0.5
0
10
20
30
40
50
87
Rb/86Sr
1.9
1.7
Biotite
1.3
87
Sr/86Sr
1.5
1.1
Feldspar
0.9
Tusas River Canyon
Age = 1509±40 Ma
Initial 87Sr/86Sr =0.731±0.031
MSWD = 20
WR
0.7
Sphene
0.5
0
10
20
30
87
40
50
60
86
Rb/ Sr
Figure 2.3. Sphene-whole rock-biotite mineral isochron of Tres Piedras Granite type
locality and Tusas River Canyon. Isochron plotted using Isoplot 3.0 (Ludwig, 2001)
83
1.9
Biotite
1.7
Biotite
1.3
87
86
Sr/ Sr
1.5
1.1
0.9
WR
0.7
Age = 1503±65 Ma
87
86
Initial Sr/ Sr =0.732±0.030
MSWD=15
WR
Sphene
0.5
0
10
20
30
87
40
50
60
86
Rb/ Sr
Figure 2.3 continued. The mineral separates (sphene, biotite) and wholerock from both Tusas River Canyon and Tres Piedras Granite type locality
sample plotted together. Isochron plotted using Isoplot 3.0 (Ludwig,
2001)
84
data-point error ellipses are 2σ
1800
0.32
1600
206
Pb/
238
U
0.28
1400
0.24
Tres Piedras Type Locality
Intercepts at
1732.7±6.7 Ma
MSWD = 0.92
1200
0.20
1000
0.16
1.5
2.5
3.5
207
Pb/
4.5
235
U
data-point error ellipses are 2σ.
0.4
1800
U
0.3
206
Pb/
238
1400
0.2
1000
Tusas River Canyon
Intercepts at
90±20 & 1729±20 Ma
MSWD = 3.5
600
0.1
200
0.0
0
1
2
3
207
Pb/
4
5
6
235
U
Figure 2.4. Concordia plot of U-Pb analysis of single zircons analyzed by LAMC-ICPMS. Concordia plotted using Isoplot 3.0 (Ludwig, 2001)
85
0.1
Rb/K
Sr/K
ratio of cps
0.01
0.001
0.0001
0.00001
0
200
400
600
800
1000
Distance in microns
Rb/Sr (cps ratio)
1000
100
10
0
200
400
600
800
1000
Distance in microns
Figure 2.5. LA- ICPMS measurement of concentration ratios across a feldspar
grain from Tres Piedras Type Locality.
86
CHAPTER THREE
Trace Element & Pb Isotope Studies Of The Kutch Volcanics Of NW-India
3.1 Introduction
The Deccan Traps were formed at the end of the Mesozoic era by outpouring of enormous lava
to form a large continental flood basalt province over vast areas (~500,000 km2) of Western,
Central, and Southern India. Compositionally the Deccan lavas are primarily tholeiitic and alkali
basalts with subordinate carbonatites that erupted from multiple centers of the Western Ghats,
India, the most prominent being a shield volcano like structure. The enormous size and a 65 Ma
(K/T boundary) eruption age of the Deccan trap magmas have made it a current and most
interesting topic of research.
The Deccan Volcanic Province (DVP) is thought to be linked to the Re’union plume, which is
responsible for the volcanic activity on Re’union Island in the Indian Ocean (Duncan, 1981;
Morgan, 1981). All current models of plate reconstruction indicate that the Indian subcontinent
drifted northward subsequent to the break up of Gondwanaland, its western margin passing over
the newly initiated Re’union hotspot at around 60–66 Ma (Karmalkar et al. 2000). Thus Deccan
volcanic province records the first magmatism from the plume head of the Re’union hotspot.
Although plume impact is a widely accepted hypothesis for the Deccan Volcanic Province
(DVP), details such as the exact location of the plume, the duration of plume-related volcanism,
interaction of the plume with the ambient lithosphere-asthenosphere and the nature of the plumeaffected mantle are still debated.
Age data on the alkaline rocks from the DVP reveal that the alkaline magmatic activity
overlapped with the main period of tholeiitic volcanism at ~66 Ma. (Pande et al., 1988).
However Basu et al., (1993) noticed that the alkaline magmatism north of the DVP is slightly
older (~3 Ma.) compared to the areas further south. The alkaline rocks in Kutch region of Gujarat
State (NW India), which lies to the NNW of the main DVP are nowhere seen to be in direct
contact with the Deccan tholeiites (Fig.3.1). This is in contrast to the minor alkaline activity on
the west-coast of India farther south, which is intrusive into the Deccan Lavas (Karmalkar et al.
2000). The occurrence of alkaline rocks in Kutch and the geographic disposition of the Kutch
lavas in relation to the main tholeiitic province, is therefore significant (Karmalkar et al., 2000,
1998). If the alkaline rocks of Kutch represents the initiation of DVP then it would also represent
the first melts generated from the Re’union plume head.
In this study, I report trace element, REE and Pb isotope data on 11 samples of alkaline and
tholeiitic basalts from Kutch region. Available Nd and Sr isotopic data on the same set of
samples (Bizimis, unpublished data) identifies three potential mantle components; Re´union
plume, continental lithosphere and asthenosphere (Indian MORB-like). This study intends to
illuminate whether or not the trace element and the Pb isotope ratios, in conjunction with the SrNd isotope ratios can identify the end member components of Kutch lava and will compare and
contrast the chemical and isotopic signatures of the Kutch volcanic rocks with the main DVP.
87
Samples of the alkali basalts from Dharam hill, close to Nakhatrana in the Kutch area, give an
age of 68.5+2 Ma which place their eruption prior to the peak activity of Deccan volcanism at 65
m.y. (Venkatesan et al., 1986). Tholeiitic basalts of the Deccan Traps that outcrop in the Kutch
area have yielded an age of 66.8+0.3 Ma (Pande et al.,1988). Thus the ages are within error of
each other. This work is a part of a larger project that proposes to precisely date the age (40Ar–
39
Ar) of Kutch volcanism in order to determine the temporal relationship with the main DVP. If
the Kutch lava marks the beginning of the Deccan eruption, it will be important to study the
chemical and isotopic character of these lavas in order to determine the early chemical characters
of the Re´union plume and it’s interaction with the subcontinental lithospheric mantle and the
asthenosphere (Indian MORB like).
3.2 Geological Setting
Recent stratigraphic reconstructions and isotopic age determinations indicate the following
sequence of Gondwana breakup: Gondwanaland breakup into eastern and western component
and departure of the India–Madagascar from Africa occurred at around 160 Ma (Plummer &
Belle, 1995). Following the initial rifting a number of mantle plumes or hotspots played
significant role in further continental breakup (Storey et al.,1995). India–Madagascar rifted from
Antarctica at around 130 Ma (Chand et al., 2001) followed by breakup between India–Seychelles
and Madagascar c. 88 Ma (Storey et al., 1995). The breakup between –Seychelles and
Madagascar is linked to the development of the Marion hotspot. Beginning of the rift between
India–Seychelles and Madagascar is dated at 96 Ma and its termination at about 84 Ma by
Plummer (1995). During this Late Cretaceous break-up event, the western margin of India
presumably rifted off the eastern margin of Madagascar; however, equivalent Cretaceous
volcanism is not well documented from western India. Possible exceptions include mafic dykes
in mainland south-west India (Radhakrishna et al., 1994, 1999), and the acid volcanic rocks of
the St. Mary islands (Valsangkar et al., 1981) that form a chain of small islands trending NW–
SE, off the western coast of central India. Naganna (1966) argued that the acid volcanic rocks
represented an early phase of Deccan flood basalt province (expected age of crystallization being
slightly greater than 65 Ma), but whole rock K–Ar ages range between 80.3 ± 1.7 Ma to 97.6 ±
2.3 Ma (Valsangkar et al., 1981) suggest that this magmatism might be better correlated with an
older event related to India–Madagascar break-up (Fig.3.1).
Next development in the sequence was rifting of Seychelles from India. Seychelles Plateau
consists of a sliver of continental crust that is traced back to some 700 Ma. (Plummer &
Belle,1995). Presumably Re’union hotspot was the cause of separation at about 65 Ma at the
Cretaceous/Tertiary (K/T) boundary. Gnos et al. (1997) corroborate the timing of separation and
reported a counterclockwise movement of India at about the same time. This was the
approximate time of Deccan volcanic event, development of the Carlsberg Spreading Ridge and
initial opening of the north-western Indian Ocean basin (Plummer & Belle, 1995).
It is postulated that the Deccan volcanic sequence in India was a consequence of the passage of
the northerly drifting Indian subcontinent over the Re’union starting plume in the Late
Cretaceous (Morgan, 1981). This plume also produced the three-rift Cambay triple junction, the
three arms being the West Coast graben belt (also known as Kachehh Rift), the Narmada-Tapi
rift zone and the Cambay rift (Fig.3.1).
88
The Kutch region is situated to the northwest of the main Deccan Volcanic province and
proximal to the Cambay Triple junction (Fig.3.1). It lies to the south west of the zone of low
seismic velocities which is interpreted as fossil plume head of the Deccan Plume (Kennett and
Widiyantoro, 1999). The tectonic evolution of the Kutch region has been attributed to the rifting
process and the region accordingly marks the site of paleo-rift graben whose evolutionary history
dates back to the Mesozoic times (Biswas, 1987).
Consistent with the northward passage of Greater India over the Reunion hotspot in the latest
Cretaceous, the geochemical stratigraphy of the south- western Deccan exhibits a southwardyounging progression of volcanism (Beane et al. 1986, Devey et al., 1986). The stratigraphic
relationship of flows in the little-studied northwestern part of the province to those in the
southwest is unknown, but at least some of the former may represent stratigraphically lower
(earlier) levels of the overall volcanic succession (Mahoney et al., 1985). An important
difference between the few northwestern areas that have been studied and the southwestern
Deccan formations is that alkalic and transitional flows are common in the northwestern sections,
whereas the southwestern lavas are almost exclusively tholeiitic basalts (Mahoney et al., 1985).
In order to ascertain the exact relation of the Kutch volcanics with the main DVP precise dates
are required. Without obvious stratigraphic relationship of the Kutch volcanics with the main
DVP it is important to chronologically place the Kutch volcanics in the broader picture. If they
are slightly older than the Deccan volcanics they might represent the earlier phase of volcanism
from the Re’union hotspot; if the age of the volcanics date back to the Jurassic the Kutch
volcanism might be related with the Marion hotspot. This work will analyze the trace element
and Pb isotope data in conjunction with the available Sr-Nd data on the 11 alkaline and tholeiitic
samples from Kutch to chemically compare them with the main DVP and identify their
affiliation.
3.3 Deccan Stratigraphy
Recent stratigraphic classification of the western Deccan (referred as the main DVP or the
southwestern Deccan in this study) between the Igatpuri and Amboli area (hatched area in Fig.
3.1) has divided the lava sequence to eleven formations with a maximum combined thickness
exceeding 3 km, based on both field markers and general geochemical similarity of the flows
(Cox and Hawkesworth, 1985; Devey and Lightfoot, 1986). The basalts lie upon and are largely
surrounded by Precambrian continental crust, much of it Archean. 40Ar–39Ar dating of samples
covering most of the exposed stratigraphy in the western Deccan indicates an age of ~66 Ma
with no measurable difference between the upper and lower part of the succession, implying <2
Ma period of eruption for the major phase of Deccan volcanism (Duncan and Pyle, 1988). Table
3.1 summarizes the stratigraphic nomenclature and thickness of the southwestern Deccan basalt
formations.
89
3.4 Previous Work – Sr and Nd Isotopic Data
Sr and Nd isotopes have been measured on 7 samples of alkali basalts with olivine phenocrysts
and 4 tholeiites from the Kutch region (Bizimis, personal comm.). Figure 3.2 compares the
initial Sr and Nd isotopes of the Kutch lavas with the main DVP basalts. Sr and Nd isotope data
for the main formations of the southwestern Deccan from Cox & Hawkesworth, 1985, Lightfoot
& Hawkesworth, 1988, Peng et al.,1994, Peng et al., 1998 are plotted in Figure 3.2. Ambenali
basalts with the highest initial 143Nd/144Nd (0.512640 to 0.512914) and the lowest (87Sr/86Sr)I
(0.7038 to 0.7060) has been identified as the least contaminated with continental material. Nd
and Sr isotopic data of the 10 formations (all except Desur of Table 3.1) define two general
trends; the Ambenali-Poladpur-Bushe trend that also includes lavas from Khandala,
Bhimasankar, Thakurvadi, Neral and Igatpuri –Jawhar and Ambenali-Panhala-Mahabaleshwar
trend (not all the formations are shown in Fig.3.2). Lavas of the Bushe formation are
characterized by very high (87Sr/86Sr)I (0.7134 to 0.7202) and very low 143Nd/144Nd initial
(0.511889 to 0.511684). They appear to be contaminated by an Archean, broadly granitic crustal
end-member (Peng et al., 1994). The Poladpur, Khandala, Bhimasankar, Thakurvadi, Neral and
Igatpuri –Jawhar lavas define a Sr-Nd array lying between those of Ambenali and Bushe
formations and appear to have been contaminated to intermediate degrees. The Mahabaleshwar
formation, directly below Ambenali exhibits an isotopic array that overlaps that of the Ambenali
but extends to low 143Nd/144Nd (0.512816 to 0.512258) and possesses much lower (87Sr/86Sr)I
(0.7040-0.7055) than the Bushe. The Ambenali-Mahabaleshwar trend has been proposed to
reflect the mixing of isotopically Ambenali like magmas with one of the following components:
a) ancient granulite crust (Mahoney, 1988; Lightfoot et al., 1990) or b) old, stabilized continental
lithospheric mantle (Lightfoot & Hawkesworth, 1988; Lightfoot et al., 1990). The question of
which of these components was involved is still the subject of investigation.
In summary, most of the formations possess Nd-Sr isotopic compositions reflecting significant
continental lithospheric influences in their origin with Bushe reflecting the highest degree of
contamination. The main exception is the 500 m thick Ambenali Formation in the upper part of
the sequence, which is relatively uncontaminated. However, the Ambenali basalts are not
equivalent to modern oceanic island products of the Re’union hotspot. Instead, the trace element
characteristics of the least-contaminated Ambenali lavas resemble those of transitional MORB,
and their isotopic signatures have been interpreted as a mixture of modern Re’union Island-type,
broadly Indian MORB-like, and minor continental lithospheric components (Mahoney, 1988).
The isotope results obtained from the least differentiated, silica-undersaturated samples from the
complexes of Barmer and Bhuj in the northeastern Deccan (Simonetti et al., 1998) are plotted in
Figure 3.2 for comparison with the main DVP lavas. The samples from Bhuj fall between the
fields defined for present-day Indian MORB and that representing the composition of Re´union
‘plume’ mantle at ~65 Ma. The melilitites from Barmer contain higher radiogenic Sr and lower
Nd isotope values than the samples from Bhuj. In addition, the melilitites from Barmer contain
similar initial 87Sr/86Sr ratios but slightly lower initial 143Nd/144Nd ratios than the composition of
the Re´union plume. Picritic and basaltic lavas from three drillholes in the northwestern Deccan
Traps, near Kutch area (Peng & Mahoney, 1995) are plotted to compare the Nd-Sr isotopes with
the main Deccan lavas. The age-corrected 143Nd/144Nd values of the drillhole lavas vary from
90
0.512855 to 0.512419, and 87Sr/86Sr, varies from 0.70414 to 0.70784; these ranges, although
substantial, are smaller than observed in the southwestern Deccan.
In the age corrected Sr-Nd space the Kutch lavas show three groups (Fig. 3.2). The tholeiites
follow the Ambenali-Poladpur-Bushe trend with low initial 143Nd/144Nd (0.512377 to 0.512621)
and high (87Sr/86Sr)I (0.70566 to 0.70894). Five of the analyzed alkali basalts plot in very close
cluster (143Nd/144Nd) i varying between 0.512839 to 0.512885 and (87Sr/86Sr) i ranging from
0.70363 to 0.70379 between the fields defined by CIR and Re’union plume at 65 M.a. Two of
the alkali basalt samples plot generally on the Ambenali-Mahabaleshwar trend but they have
lower (87Sr/86Sr)I than the Mahabaleshwar samples and might represent a different source than
the rest of the alkali basalts.
3.5 Analytical Technique
Eleven samples from the Kutch region have been analyzed for their trace elemental composition
and lead isotope, using ICPMS and TIMS (MAT262) respectively and the results are shown in
Table 3.2.
Trace elements for the alkali basalts/tholeiites were determined by solution ICP-MS analysis
using a ThermoFinnigan Element ICPMS. Small chips of each sample were handpicked and
crushed with agate mortar and pestle. 30 mg of each sample was weighed into screw top Teflon
beakers and dissolved with 2 ml of 3:1 distilled HF-HNO3. After drying on a hot plate at 120oC,
the samples were allowed to reflux overnight in concentrated HNO3. Samples were then dried
again and brought to a final volume of 60 ml with 2% HNO3. The sample solutions were further
diluted in the ICP vials for target concentration of 100 ppm TDS. 1 ppb of Indium was added as
internal standards and drift corrections for each analyzed mass were applied by interpolating
with the internal standard. The solution ICPMS analyses were calibrated against a single solution
of well-characterized Hawaiian basalt standard BHVO-1 prepared identically to the samples.
BCR-1 was also prepared like sample and calibrated against the BHVO-1 to check the precession
of measurement.
Pb isotopic composition was determined in the Finnigan MAT-262. 120 mg of sample powder
was leached for ~20 min in cold 6 N HCl in an ultrasonic bath, rinsed several times with Quartz
Distilled (QD) water and dissolved in 5 ml of 3-1 distilled HF-HNO3. After drying on a hot plate
at 120oC, the samples were allowed to reflux overnight in concentrated HNO3 for 3 times, dried
and fluxed again in 5 ml of distilled 6 N HCl twice and finally dried in concentrated HBr twice.
Pb was separated using a two-column procedure (Abouchami et al., 1999) and run on a single
commercial Re filament using MAT 262. Total blanks were better than 0.3 ng for Pb; and
repeated (n=18) analysis of isotopic standard NBS 981 yielded 206Pb/204Pb = 16.90 ± 0.02,
207
Pb/204Pb = 15.45 ± 0.02, 208Pb/204Pb = 36.60 ± 0.04.
3.6 Results
The trace element, REE and Pb isotope data are shown in Table 3.2. Concentrations of selected
trace elements of the Kutch samples arranged in order of ascending compatibility and normalized
to primitive-mantle and CI chondrite concentrations are shown in Fig 3.3a and 3.3b respectively.
91
The samples can broadly be divided into two groups, the alkali basalts with olivine phenocrysts
and the tholeiites.
3.6.1 Alkali Basalts
Alkali basalts are characterized by an overall smooth trace element pattern with decrease in
absolute abundance from Nb to Yb (Fig. 3.3a) and a small range in HREE and Y contents (Fig.
3.3b). Compared to MORB concentrations these samples show enrichment in Nb and LILE. All
are enriched in LREE. The ratios of Tb/Yb vary between 0.59 and 0.87 compared to the value of
0.18 in MORB and suggests fractionation of HREE. The ratios of Ti/Zr for the Kutch alkali
basalts vary between 61.5 and 104.6 whereas the Zr/Y varies from 8.4 to 13.2 and Ti/Y ranges
from 582 to 1015. For most primitive “MORB”, the ratios Ti/Zr, Zr/Y, and Ti/Y are about 100.5,
1.95 and 196 respectively (Salters and Stracke, 2004). The Indian Ocean MORB however is
reported to have a mean Ti/Zr ratio of 84 and therefore is an exception to the values given for NMORB. The (La/Sm)n ratio for the Kutch samples is in the range of 2.7 – 4.7, which is higher
than the range of 0.86– 1.5 given for MORB. The Lu/Hf values for the Kutch samples range
from 0.0299 to 0.0523, which are much lower than for MORB-source mantle (~0.3). All the
Kutch alkali basalts show positive Ba and Sr anomalies and a prominent negative Yb anomaly.
Overall, the LREE are variously enriched over HREE ((La/Yb)n=21.7–57.3). No Europium
anomaly is observed in the samples, and the scatter in trace element abundances maybe regarded
as primary, or as reflecting crystal fractionation involving predominantly olivine and
clinopyroxene (Simonetti et al., 1998). Due to the lack of Eu anomaly plagioclase is not
considered to be involved in fractional crystallization. Fisk et al. (1988) have reported similar
trends for the rocks of Reunion Island, which they have related to melt, differentiation.
The incompatible element patterns of the alkaline rocks from Kutch region are broadly consistent
with a dominant OIB type source component. All of the alkali basalts have high abundances of
Nb, with low Zr/Nb (3.0–5.1) typical of OIB and contrary to the higher values reported for
MORB (6–10). The Ce/Pb (23.5-26.3) ratios for the Kutch alkaline rocks are consistent with the
relatively constant values for these ratios reported in OIB (25±5) (Sun and McDonough, 1989).
The isotopic data for some samples from Kutch area characterized by lower initial 87Sr/86 Sr
(0.70357–0.70396) and higher initial 143 Nd/144 Nd (0.512839–0.512885) ratios. These ratios
closely correspond to those for Reunion.
Most of the Kutch alkali basalts show a narrow range of (La/Yb)n (21.7–28.0)(except the two
most enriched samples) and (Tb/Yb)n ratios (2.6–3.8) but have low and uniform HREE ((Yb)n
contents 0.52 to 0.82). This is consistent with partial melting of variable degrees with garnet
being present in residual phase. A garnet signature is also evident from the conjunction of
fractionated HREE and low Y, Sc and HREE contents in magmas with Ni >150 ppm. Lu/Hf data
may also be used to place some constrains on the source mineralogies. This is because the
partitioning systematics between melt and residual mantle for Lu and Hf are strongly affected by
the presence of garnet. The Lu/Hf values for the Kutch samples range from 0.030 to 0.052, these
lower values imply melting in the garnet stability zone.
92
3.6.2 Tholeiites
Considering the continental setting in which the Kutch samples have been emplaced, crustal
contamination is a distinct possibility. In Fig. 3.3a, the average composition of the continental
crust has been plotted along with the tholeiite samples. They are characterized by a LREE pattern
similar to that of the alkali basalts, but flatter HREE (Fig. 3.3b). They show a pronounced
positive Pb anomaly and only slightly positive La anomaly like that of the average continental
crust. Negative Nb anomaly is similar to the continental crust but they show only modest
negative Ti anomaly unlike that of the continental crust. (Th/Nb)n ranges between 5 and 7 as
against 6.18 in continental crust (Rudnick and Gao, 2003). The alkali basalt samples on the
contrary are characterized by lack of Nb and Pb anomalies, lower HREE, and (La/Sm)n higher
than the bulk continental crust. The La/Nb ratios of the Kutch tholeiites (1.32 to 1.57) are
comparable with those of the continental crust (~1.5). The indices of crustal contamination such
as Ce/Pb or Rb/Sr ratio for the Kutch tholeiits vary from 6.86 to 12.09 and 0.08 to 0.15,
respectively, as against values of 4.13 and 0.12 for the continental crust (Rudnick and Gao,
2003). These similarities with the bulk continental crusts are considered to have been produced
by significant crustal contamination.
3.7 Comparison with the DVP
Figure 3.4 illustrates Ce/Pb vs. La/Nb and La/Sm vs. Sm/Yb in the Kutch rocks. For comparison
the CIR, Re’union, southwestern Deccan lavas (or lavas from the main DVP) and Northwestern
lava fields are shown. Ce - Pb and La - Nb have similar partition coefficient during mantle
melting and any difference in the ratio will illustrate the source character rather than degree of
melting and/or fractional crystallization. On the contrary the La/Sm vs. Sm/Yb defines the entire
slope of the REE and as the elements have different partition coefficient a difference in the ratio
will be due to different degrees of melting and/or fractional crystallization.
In the Ce/Pb vs. La/Nb plot the tholeiites shows a crustal signature with high La/Nb and low
Ce/Pb. The alkali basalts plot in the Re’union field. For comparison the drill hole lavas from
northwestern Deccan and lavas from the main DVP are plotted also. They show a wide variation,
ranging from Re’union type source to extreme crustal contamination. In the La/Sm vs. Sm/Yb
plot the alkali basalts plot above the Re’union field with the two most enriched alkali basalts
having the highest La/Sm and Sm/Yb ratio. The DVP and drillhole samples plot at a much lower
Sm/Yb ratio (<3.5) though they display a wide range of La/Sm ratio (1.7 to 8.5). The tholeiites
plot on the DVP field.
A plot of present day 206Pb/204Pb vs. 207Pb/204Pb (Fig. 3.5a) shows the data listed in Table 3.2, in
addition to the fields for Archean basement, Indian MORB and the Re´union plume component.
As in the case for their initial Nd and Sr isotope data (Fig. 3.2), the 5 alkali basalt samples from
Kutch contain Pb isotope ratios that plot on and between the Re´union and Indian MORB fields
(Fig. 3.5a). One of the two most enriched alkali basalt that follow the Ambenali-Mahableshwar
trend in the initial Nd-Sr isotope space falls in between the Re´union and Indian MORB fields;
the other one falls slightly below the array. The Pb isotope ratios for tholeiite samples from
Kutch are clearly more radiogenic compared with the remaining samples. The tholeiite samples
are enriched in 207Pb except for BH-2 which is enriched both in 207Pb and 206 Pb. When compared
93
with the southwestern Deccan lavas the alkali basalts overlap with the Ambenali in the
206
Pb/204Pb vs. 207Pb/204Pb space. A (87Sr/86Sr)I vs. 207Pb/204Pb plot of the southwestern Deccan
shows two trends diverging from the CIR-Re’union overlapping field (Fig. 3.5b). 1.
Mahabaleshwar trend with increasing (87Sr/86Sr)I with progressively decreasing 207Pb/204Pb 2.
crustally contaminated trend with increasing (87Sr/86Sr)I and 207Pb/204Pb. The alkali basalts plot at
the intersection of the two trends on the Re’union field where as the tholeiites fall on the crustal
contamination trend.
3.8 Discussion
In the La/Sm vs. Sm/Yb plot the alkali basalt samples plot at a higher ratio than the Re’union
plume indicative of a very low degree of partial melting of Re’union type source. The primitive
mantle normalized ratios of (Sm/Yb)n in the Kutch rocks are 4–6. This type of values are
observed when melting initiates in presence of residual garnet (Ellam, 1992). Here we present a
batch-melting model and incremental batch-melting model of garnet peridotite in the garnet
stability field and spinel stability field of Re’union type source. Since the Re’union plume is not
directly accessible, its composition has to be inferred via constraints derived from rocks, whose
origin and composition are related to the plume related magmatism, like the Deccan volcanics.
The isotopic composition of the 5 alkali basalt lavas that are thought to be produced by partial
melting of the Re’union plume is used to estimate parent-daughter ratios (e.g., Sm/Nd, Rb/Sr,
Lu/Hf) and the respective element concentrations in the Re’union plume (e.g., Rb, Sr, Sm, Nd,
Lu, Hf, etc).
The age corrected isotopic composition of alkali basalts of the Kutch directly reflects the isotopic
composition of its source, the Re’union plume at 65 Ma, and is the most reliable constraint on
the trace element ratio formed by the parent and daughter element of the isotopic system
considered. For example, by measuring 143Nd/144Nd in alkali basalts, the ratio of the parent
element Sm and the daughter element Nd (Sm/Nd) can be calculated with some information on
the isotopic evolution of the Re’union plume. These calculated parent-daughter ratios form the
framework for the estimates of other trace element ratios.
The present day isotopic ratios are used to calculate the elemental ratios of the Re’union plume
assuming the average age of the Depleted Mantle to be 2.2 Ga (Chase and Patchett, 1988;
Condie, 2000) is given in Table 3.3.
Table 3.3. Estimating the source composition of the Kutch alkali lavas.
Present day
Reservoir ratio.
At 65 Ma source
ratio
(87Sr/86Sr)
(87Rb/86Sr)
(143Nd/144Nd) (147Sm/144Nd)
0.7045
0.082182
0.512638
0.703597
0.0526
0.512867
(176Hf/177Hf)
(176Lu/177Hf)
0.1967
0.282772
0.0332
0.2213
0.2831094
0.0432
Present day (87Rb/86Sr) calculated assuming (87Sr/86Sr)BABI of 0.69989 at 4.57 Ga. (87Sr/86Sr)65,
(143Nd/144Nd)65, (176Hf/177Hf)65 (Bizimis, personal communication) ratios are age corrected
average of 5 alkali basalt samples measured for Kutch volcanics. Resulting elemental ratios are
(Rb/Sr)65 = 0.01817 (Sm/Nd)65 = 0.36608 and (Lu/Hf)65 = 0.3039.
94
Figure 1 from Salters and Stracke (2004) is used for estimates of the individual elements and the
inter-relationship between the estimates of the different elements. Lu concentration has been
choosen as the ‘‘anchoring point’’ of the trace element pattern and fixed at Bulk Silicate Earth
(0.0675 ppm) of McDonough and Sun, (1995). With the Lu concentration as a starting
concentration, the Lu/Hf ratio derived from the Hf-isotope systematics is used to estimate the Hf
concentration. The Hf/Sm ratio of the Re’union is assumed to be chondritic and thus the Sm
concentration of the Re’union plume is calculated from the Hf concentration. Nd concentration is
calculated from the Sm/Nd ration of Re’union at 65 Ma. Again Nd/Sr ratio of the Re’union is
assumed to be chondritic that gives the Sr concentration and hence Rb concentration. Thus after
Lu, Hf, Sm, Nd, Sr and Rb are anchored the elements in between are assigned values to generate
a smooth curve in a Primitive Mantle normalized spider diagram. Thus the La, Sm and Yb
concentrations are determined to be 0.35, 0.317 and 0.42 ppm respectively.
In a Primitive Mantle normalized plot of (La/Sm)n vs. (Sm/Yb)n (Fig. 3.6a and 3.6b)
continuous lines are for melting of Fertile Lherzolitic mantle. The Kutch alkali basalts plot can
be produced by approximately 1.6% batch melting of a garnet peridotite source with 53%
olivine, 8% orthopyroxene, 34% clinopyroxene and 5% garnet. Incremental batch melting of the
same source yield similar results of approximately 1.6 to 1.8% melting of the source. The two
extremely enriched alkali basalts cannot be produced with similar source and possible was
derived from a different source.
For the tholeiites, normalize incompatible element patterns of contaminated lavas should reflect
the patterns of both their source and contaminant (whether a bulk rock, partial melt, etc.)
weighted by the relative contribution of each. Patterns for the Kutch tholeiites show several
important features typical of continental crust with negative Nb spikes and strongly positive Pb
peaks and either they lack negative Ti spikes or have modest negative Ti anomalies. Published
analyses of Precambrian crustal rocks (e.g., Weaver and Tarney, 1980, 1981 ; Thompson et al.,
1983; Volpe and MacDougall, 1990) indicate that high-SiO2 types typically have very prominent
negative Ti spikes in their incompatible element patterns, as well as marked enrichment in the
highly incompatible elements other than Nb. Many basic amphibolites and basic to intermediate
granulites are enriched in the highly incompatible elements and show only slight negative Ti
spikes or lack them altogether (e.g. Weaver et aI., 1977; Khandelwal and Pandaya, 1988;
Gopalan et al., 1990). For example, Weaver et al. (1977) identified four groups of Archean
granulites in the Madras area of southern India: basic granulites, intermediate granulites,
chanockites, and khondalites. Incompatible element patterns for these rocks show that the basic
and intermediate varieties possess patterns qualitatively comparable to those of the Kutch
tholeiites. Patterns of the basic granulites have negative Nb spikes, a pronounced positive Pb
peak, and lack negative Ti spikes; patterns for the intermediate granulites, have higher Ba
contents.
We now apply a simple model to assess whether crustal assimilation associated with fractional
crystallization can plausibly account for the isotopic and incompatible element characteristics of
the common signature lavas (Fig. 3.7a and 3.7b). For the model the magma temperature was
assumed to be 12000C and the temperature at the base of the crust is assumed to be 8000C. We
selected the uncontaminated end-member to be similar to either CIR or a primitive Re’union type
95
parent magma. To fit the trace element patterns of the common signature lavas, we added
variable amounts of a given contaminant composition to that of the uncontaminated parental
magma. Several types of bulk rock contaminants were tried. The best value (i.e. by best visual
fits) were achieved with bulk rock mixture of 75% CIR basalt with 18% basic granulite (n=7)
and 6% charnockite (n=28) analyzed by Weaver et. al., 1978. A fractional crystallization (after
the Rayleigh’s fractional crystallization model) of 18% olivine was required to generate the
observed pattern. In absence of the major element chemistry it was difficult to pinpoint the
nature of the fractionating crystals (olivine, clinopyroxene and/or plagioclase). Olivine
fractionation gave the best visual fits. The trace element pattern for the most contaminated
tholeiite, BH-2 could be generated by mixing 75% CIR basalt with 20% charnockite and 5 %
basic granulite (n=10) and fractionating 30% olivine. In absence of the major element
concentrations of the Kutch tholeiites, a simple calculation was done to check the SiO2 content of
a contaminated MORB with 20% charnockite. Assuming 46% SiO2 concentration for CIR
basalts, 20% bulk charnockite (70% SiO2) was mixed. The resultant rock will have a 50.8% SiO2
which is still within the range of silica concentrations of the Deccan basalts. Different
proportions of granulite, charnockite and khondalites were mixed with the Re’union magma to
test the plume as a plausible end member but we failed to generate the observed Nb anomaly and
slight Ti anomaly.
Neither the specific chemical nor isotopic characteristics of the lower crustal rocks located
beneath the Deccan are known. However it is interesting that the optimum mixture of the mantle
and granulite and charnockite end-members suggested by trace element calculations also yield
appropriate Nd-Sr isotopic values. When mixed with granulites and charnockites from
Krishnagiri, South India (west of Madras Granulites) the isotopic composition of the tholeiites
could be generated by 18-24% mixing of the Krishnagiri granulites and charnockites analyzed by
Peucat et.al., 1989 (Fig. 3.7c).
Lead isotope and uranium abundance data available for the Krishnagiri charnockites and
granulites have much lower 207Pb/204Pb and 206Pb/204Pb ratios than it is observed for the Kutch
volcanics. However Chakrabarti and Basu (2006) measured some impact breccias from Lonar
Crator that have lower 206Pb/204Pb and 208Pb/204Pb ratios but higher 207Pb/204Pb ratios compared
to the host basalts. They argued that a major component of the Lonar impact breccias was
derived from melting of Archean basement rocks. The basement beneath the Lonar region is
believed to be similar to the Dharwar craton of peninsular India. Based on their similar Pbisotopic compositions with the breccia rocks, Chakrabarti an Basu suggested that Archean
Chitradurga Group of rocks of this craton to be present in the basement beneath the Deccan lavas
of the Lonar region. Thus the lead isotopic composition of the Kutch tholeiites could have been
modified by the contamination of the Archean basement rocks.
3.9 Conclusion
The Kutch lava shows three types of end members, OIB - type end member, Indian MORB and
Archean crust. The trace element pattern and Sr-Nd-Pb isotope ratios of the alkali basalts are
similar to that of OIBs and in absence of precise age data it is difficult to pinpoint it’s source
(Re’union or Marion hotspot). Elemental data indicate that these lavas are probably the products
of comparatively small degrees of partial melting, less than for the southwestern Deccan lavas. If
96
the Re’union plume is responsible for the Kutch lavas that erupted early in the sequence, gradual
increase with time in the degree of partial melting in the Deccan (Beane et al. 1986, Devey et al.,
1986) is likely to be at least partly responsible for regional north-to-south decreases in average
Nb/Zr, Nb/Y. The two most enriched samples have lower (143Nd/144Nd)I than the rest of the
alkali basalts. The isotopic ratios of those two samples can be explained by mixing of
charnockites with either Re’union or Indian MORB type end members but that fails to explain
the trace element pattern. The two enriched samples might have a different source and/or genesis
than the rest of the alkali basalt and needs to be examined further in details.
The tholeiites from the Kutch are significantly contaminated by crustal material of granulitic and
charnockitic composition. The trace element pattern could be explained by mixing Indian MORB
with the Archean crust from South India. Assuming Re’union as the uncontaminated endmember source failed to produce the observed Nb anomaly in the tholeiites. The Sr-Nd isotope
ratios were compatible with our model too. Thus if Kutch lavas are the first phase of eruption of
the DVP then it’s important to note that both Re’union and Indian MORB played significant role
during the early phase of Deccan eruption.
Table 3.1. Stratigraphic nomenclature and thickness of the
southwestern Deccan formations (modified after Peng et al., 1994).
Subgroup
Formation
Max. Thickness
(87Sr/86Sr)I
WAI
Desur
~100 m
0.7072 – 0.7080
Panhala
>175 m
0.7046 – 0.7055
Mahabaleshwar
280 m
0.7040 – 0.7055
Ambenali
500 m
0.7038 – 0.7060
Poladpur
375 m
0.7053 – 0.7083
Bushe
325 m
0.7134 – 0.7202
Khandala
140 m
0.7068 - 07107
Bhimashankar
140 m
0.7067 – 0.7076
Thakurvadi
650 m
0.7066 – 0.7112
Neral
100 m
0.7082 - 0.7104
Igatpuri-Jawar
>700 m
0.7089 - 0.7124
LONAVALA
KALSUBAI
97
Table 3.2. Trace element and Pb-isotope results.
Sample
BH 19.3 BH 12.3 BH 13.1 BH 16.2
Alkali
Alkali
*Alkali *Alkali
Basalt
Basalt
Basalt
Basalt
206
19.02
18.83
18.37
Pb/204Pb 18.58
207
15.53
15.56
15.50
Pb/204Pb 15.55
208
40.18
39.07
38.37
Pb/204Pb 39.14
8.86
11.95
5.58
8.35
Li
55.22
103.24
27.70
25.51
Rb
877
1516
606
572
Sr
25.75
27.74
19.38
20.82
Y
290
366
188
178
Zr
95.97
118.50
40.62
34.34
Nb
3.42
2.32
0.59
1.13
Cs
1251
2321
470
418
Ba
65.41
87.11
28.91
27.54
La
122.73
157.85
61.13
58.34
Ce
13.82
17.49
7.52
7.02
Pr
55.29
67.03
31.76
29.88
Nd
10.40
11.61
6.63
6.31
Sm
3.65
4.35
2.29
2.19
Eu
15.45
17.89
8.36
7.78
Gd
1.48
1.62
0.96
0.92
Tb
5.98
6.09
4.30
4.40
Dy
0.97
1.00
0.72
0.76
Ho
2.30
2.37
1.77
1.94
Er
1.79
1.86
1.41
1.55
Yb
0.23
0.23
0.18
0.21
Lu
6.68
7.57
4.27
4.10
Hf
5.58
6.32
2.63
2.15
Ta
4.77
5.98
2.48
2.44
Pb
9.51
11.01
3.68
3.32
Th
2.07
2.41
0.82
0.73
U
21.11
19.72
22.73
26.28
Sc
22875
22478
19676
16223
Ti
243.1
224.9
307.1
298.8
V
522.6
334.0
514.7
668.6
Cr
74.9
67.7
58.6
63.3
Co
516.1
373.9
251.6
323.5
Ni
52.0
59.8
73.8
74.7
Cu
130.9
130.7
104.4
100.8
Zn
BH 7.1
Alkali
Basalt
18.94
15.55
39.02
7.15
22.27
774
33.21
280
57.04
0.64
514
44.76
95.20
11.58
48.68
10.02
3.36
13.19
1.50
7.01
1.23
3.08
2.51
0.33
6.22
3.41
3.69
5.76
1.42
32.19
19336
298.1
510.5
63.6
270.1
58.8
117.1
* two most enriched alkali basalt samples.
98
N 1.1
Alkali
Basalt
18.64
15.53
38.76
6.37
42.09
765
22.37
206
43.53
0.64
478
36.55
72.62
8.61
36.11
7.47
2.57
9.92
1.07
4.99
0.83
2.09
1.63
0.22
4.84
2.86
2.95
5.08
1.12
27.50
16581
306.9
712.1
73.7
380.7
75.8
107.5
BH 18.1 BH 17.1 BH 2
BH 3.1 BH 3.2
Alkali Tholeiite Tholeiite Tholeiite Tholeiite
Basalt
18.68
18.71
20.09
18.31
18.34
15.52
15.62
15.83
15.61
15.66
38.75
39.06
41.10
38.75
38.90
6.68
5.93
9.30
6.72
6.20
29.41
10.82
37.32
22.50
20.58
694
141
247
207
187
21.46
26.90
45.27
40.39
36.35
220
77
262
153
133
49.17
4.69
24.62
14.18
12.63
0.63
0.50
0.60
0.61
0.56
530
120
337
326
298
36.42
7.39
34.34
19.59
16.72
74.91
15.85
72.41
39.05
33.94
8.97
2.07
8.64
4.68
4.20
37.40
9.56
35.61
19.27
17.59
7.64
2.71
8.04
4.45
4.08
2.65
0.95
2.45
1.55
1.45
10.67
2.93
9.46
5.58
5.38
1.13
0.57
1.31
0.88
0.82
4.77
4.34
8.12
6.32
5.80
0.80
0.94
1.57
1.36
1.25
1.95
2.88
4.45
4.21
3.89
1.59
2.48
3.60
3.69
3.49
0.20
0.38
0.52
0.57
0.53
4.97
1.99
6.22
3.61
3.20
3.24
0.31
1.30
0.88
0.78
3.19
1.96
5.99
5.43
4.95
5.02
1.69
7.77
3.84
3.26
1.16
0.35
1.70
0.79
0.70
23.97
52.46
34.45
43.50
40.21
16649
5784
17381
9330
8705
296.6
362.8
473.1
389.0
359.6
750.6
282.6
125.1
6.7
5.7
80.5
55.0
41.6
48.2
45.8
554.4
116.3
54.7
28.7
26.1
84.7
151.4
219.4
204.6
190.2
125.3
99.2
127.4
119.2
112.5
Kutch Volcanics
- Study Area
Main DVP
Figure.3.1. Tectonic and structural features of southern Asia and the Indian Ocean basin
(based on Mahoney et al., 2002 and Sheth, 2005). Abbreviations for localities are: Br,
Barmer, M, Mundwara; D, Dhandhuka; B, Bombay; R, Rajahmundry; G, Goa dykes; KK,
Karnataka-Kerala dykes; SMI, St. Mary's Islands volcanics. Mahoney et al., 2002 modeled
Re’union hotspot track showing expected ages in Ma.
99
0.5131
0.5130
143Nd/144Nd
0.5129
CIR
Re'union
Bhuj
Ambenali
0.5128
0.5127
Barmer
NW-Deccan
0.5126
0.5125
Poladpur
0.5124
Igatpuri-Jawhar
0.5123
0.5122
0.702
Mahabaleshwar
Khandala
Bushe
0.704
0.706
0.708
0.710
87Sr/86Sr
Figure. 3.2. Initial Nd vs. initial Sr isotope plot of the Kutch samples (Bizimis,
per. comm.), which are compared with data for Central Indian Ridge (CIR) (bold
blue field—data taken from Cohen et al., 1980; Cohen & O’Nions, 1982; Ito et
al., 1987; Mahoney et al., 1989; Rehkaemper,1997; Nauret, 2006), Re´union
(bold blue field—data from Dupre´ & Alle`gre, 1983; Fisk et al., 1988; Fretzdorff
& Haase, 2000), Northwestern Deccan (red field—data from Peng & Mahoney,
1995; ). For the southwestern Deccan Ambenali and Mahabaleshwar data field
(green field), Poladpur, Khandala, Igatpuri-Jawhar (black field) taken from Cox &
Hawkesworth, 1985; Lightfoot & Hawkesworth, 1988; Peng et al.,1994; Peng et
al., 1998. Alkaline picrites and basaltic flows from Bhuj and Barmer is from
Simonetti et al., 1998. Kutch alkali basalts are represented by red circles, the two
most enriched samples of the alkali basalt are red squares and the tholeiites are
orange triangle.
100
3a
Sample Normalized to Primitive Mantle (McDonough & Sun, 1995)
1000.000
BH 2
BH 3-1
BH 3-2
100.000
sample/standard
BH 17-1
BH 7-1
BH 13-1
BH 16-2
10.000
BH 18-1
N 1-1
BH 12-3
BH 19-3
1.000
N-MORB
AVG C ONT.
C R US T
0.100
Rb Ba Th
3b
U Nb La Ce Pb Pr
Sr Nd Zr Sm Eu
Ti
Dy Yb Y
Lu
Sample normalized to CI chondrite (McDonough & Sun 95)
200
BH 2
BH 3-1
100
sample/standard
BH 3-2
50
BH 17-1
BH 7-1
20
BH 13-1
10
BH 16-2
BH 18-1
5
N 1-1
BH 12-3
2
BH 19-3
La
Ce
Pr
Nd
Sm
Eu
Tb
Dy
Ho
Er
Yb
Lu
Figure 3.3 a) trace element compositions of Kutch basalts normalized to the Primitive
Mantle. Red lines are the tholeiites, black lines are alkali basalt and blue lines are the
two most enriched alkali basalts. N-MORB (Salters and Stracke, 2004) and Average
Continental Crust (Rudnick and Gao, 2003) are also plotted for comparison. b) The
REE of the same set of samples are plotted on a CI normalized plot.
101
30.5
Ce/Pb
25.5
20.5
15.5
10.5
5.5
0.5
0.5
1
0.5
2.5
La/Nb
1.5
2
2.5
4.5
6.5
6.5
Sm/Yb
5.5
4.5
3.5
2.5
1.5
0.5
La/Sm
8.5
Main DVP
30. 5
25. 5
20. 5
Figure.3.4. La/Nb vs. Ce/Pb and
La/Sm vs. Sm/Yb in the Kutch
Basalts. Data source same as
Figure.3.2. Symbols for Kutch
samples same as Figure. 3.2.
102
Re'union
15. 5
10. 5
CIR
5. 5
0. 5
0. 5
1
1. 5
La / N b
2
2. 5
NW-Deccan
15.9
Bushe
15.8
Archean Basement
207Pb/204Pb
15.7
NW-Deccan
Ambenali
15.6
Re'union
15.5
CIR
15.4
15.3
Mahabaleshw ar
15.2
16
16.5
17
17.5
18
18.5
19
19.5
20
20.5
21
206Pb/204Pb
16
15.9
Bhimsankar
Igatpuri-Jaw har
15.8
207Pb/204Pb
15.7
Re'union
15.6
15.5
NW-Deccan
Ambenali
CIR
Khandala
Pahala
15.4
15.3
15.2
Mahabaleshw ar
15.1
15
0.702
0.704
0.706
0.708
0.710
0.712
87Sr/86Sr
Figure. 3.5. Pb and Sr isotopic variation in Kutch Basalts. Symbols same as Figure. 3.2.
Fields drawn from same data set as Figure. 3.2.
103
8
Garnet
Stability Field
(Sm /Y b )n
6
Spinel Stability
Field
4
0.1%
0.6%
1.6%
2
0.1%
0.6%
1.6%
0
0
1
2
3
4
5
6
7
8
(La/Sm)n
Figure.3.6a. Batch-melting model of garnet peridotite and spinel peridotite.
Continuous lines are for percent melting of Fertile Lherzolitic mantle. Mineral
Fraction, Partition Coeficient and Melt modes from Salters and Stracke, 2004.
Symbols same as Figure 3.2.
(Sm /Yb)n
8
6
4
0.5%
1%
2%
5%
2
0
0
1
2
3
4
5
6
(La/Sm)n
Figure.3.6b. Incremental Batch-melting model of garnet peridotite. Continuous
lines are for percent melting of Fertile Lherzolitic mantle. Mineral Fraction,
Partition Coeficient and Melt modes from Salters and Stracke, 2004. Symbols
same as Figure 3.2.
104
100.000
BH 3.1
3.7a
BH 3.2
BH 17.1
Mixing Line
10.000
1.000
0.100
Rb
Ba
Th
Nb
La
Ce
Pb
Sr
Zr
Ti
Y
100.000
BH-2
Mixing Line
3.7b
10.000
1.000
Rb
Ba
Th
Nb
La
Ce
Pb
Sr
Zr
Ti
Y
Figure. 3.7. Primitive Mantle normalized incompatible element diagram of the Kutch tholeiites
compared with mixing different proportion of MORB with charnockites and granulites.
a) the pattern is best explained by assimilating 18% basic granulite and 6% charnockite with
75% Indian MORB and fractionating 18% olivine. b) 20% charnockite and 5% granulite is
mixed with Indian MORB and 30 % olivine is fractionally crystallized to explain the pattern for
the most contaminated sample. Trace element
105 concentration of the granulites and charnockites
taken from Weaver, 1978.
0.5135
CIR
0.513
Re'union
0.5125
17 %
24%
Granulite &
Charnockite
143Nd/144Nd
21 %
0.512
Granite, Tonalite
0.5115
0.511
0.5105
0.51
0.702
0.704
0.706
0.708
0.71
0.712
0.714
0.716
0.718
0.72
87Sr/86Sr
Figure. 3.7c. The range of Sr and Nd isotopic composition (at 65 m.y.) is shown for
the charnockite, granulite, granite and tonalite from Southern India (Peucat, 1989).
The observed isotopic ratios for the Kutch tholeiites can be generated my mixing
17-24% of granulite and charnockite of varying isotopic ratios. Symbols same as
Figure. 3.2.
106
APPENDIX
A. Analytical Techniques.
A.1. Sr and Nd chemical separation and mass spectrometry
For Nd and Sr isotopic analysis ca. 50-60 mg of sample powder was weighed in a
Savillex screw top beaker and dissolved in 3:1 HF:HNO3 mixture for about 48 hours at
about 1000C. To assure complete dissolution of phases like zircon, apatite, sphene etc the
samples were dissolved in HF/ HNO3 mixture (3:1) in high-temperature and -pressure
Teflon steel-jacketed (Parr) bombs for 48 hrs at 1800C. The bombs are left to cool down
in the next 48 hours before opening. The samples are then dried on the hot plate for 24
hrs at 1200C to remove SiF4 and H2O that is produced by the reaction of SiO2 with HF
(4HF + SiO2 = SiF4 + H2O). To remove the last drop of acid the temperature is raised to
1800C to 2000C for several hours. After complete evaporation of SiF4 the sample stopped
fuming. When the beaker is completely free from condensing acid droplets, the
temperature is lowered to 1250C and 5-6 ml of 6N HCl is added to the dry sample cake,
re-dissolved and left to dry. This step is repeated 3 times to ensure the removal of
remaining fluorides.
In the next step, the sample is re-dissolved in 0.5 ml of 2.5N HCl. The solution with
some of the un-dissolved residue is transferred into centrifuge tube and centrifuged for 6
– 9 minutes till a clear solution is achieved. This solution is now ready to load on the
cation columns to separate Sr from the REE’s.
Half of the sample solution (0.25 ml) in 2.5N HCl is loaded on quartz glass cation
columns filled with approximately 3.7 cm3cation resin. The bulk sample is eluted with
approximately 18 ml of 2.5 N HCl. Sr is collected in 5 ml of 2.5N of HCl. After the Sr is
collected 4 ml of 6N HCl is passed through the columns and the rare earth elements are
collected in another 6 ml 6N HCl.
Resin material may leak through the columns. To ensure complete decomposition of resin
material Sr fractions are dried down with 3 drops of HClO4 twice. Then 3 drops of
concentrated nitric acid is added and dried to ensure complete nitrate conversion. The
process is repeated thrice. The Sr fraction is loaded on single W filaments as a tiny drop
of 1 microliter 0.25N HNO3 mixed with equal amounts of the TAPH (10g H2O + 0.5g
TaCl5 + 3g 0.1N H3PO4 + 0.5g conc. HF) solution activator. It has been well tested that
the TAPH solution enhances ionization of Sr as well as binds it well on the filament to
retain the sample long enough for the measurement.
Sr measurements were performed on a Finnigan 262 multicollector mass spectrometer.
The measurements were performed in the dynamic mode with one peak switch. The
signal intensities were obtained after online fractionation correction based on the power
law and the 86Sr/88Sr ratio of 0.11940.
107
The signal intensities for 88Sr were always greater than 1 Volt. The granites and gneisses
had some Rb leaking into the Sr fraction even after column separation. 87Rb signal
interferes with 87Sr .Care was taken so that 85Rb signal was less than 300 counts per
second (cps), to avoid signal interferences. Corrections to the 87Sr signal required for
87
Rb interference can be neglected when 85Rb signal is less than 300 cps. In most cases
repeated “flashing” of the filament at high temperatures for a few seconds burns off the
excess Rb.
The reproducibility of the measurements is controlled by replicate measurements of the E
& A Sr isotope standard. The average measured 87Sr/86Sr ratio was 0.708006 (0.000019,
two standard deviations, n = 10). At least 10 blocks (100 ratios) were taken for each
measurement and it was ensured that the internal precision is always better than the
external reproducibility of the standard (i.e. less than 0.000019, 2 sigma error).
Nd was collected in a bulk REE fraction from the cation columns in 6N HCl. Nd and the
REE fraction was further separated on a 1.2 ml, 6 cm bed-length column of Ln resin SPS.
After collection, the Nd fraction is dried and is ready for measurement. The Nd fraction is
dissolved in ca. 1 microliter 0.25N HCl, mixed with equal amounts of 0.25N H3PO4 and
loaded as a single drop on a zone-refined Re filament. The Nd measurements were
performed using the double Re filament method.
Potential interferences on Nd are Ba, La, Pr and Ce and their oxides. They arise from
LREE and LILE enriched rocks where Nd separation was incomplete. The most common
interfering mass on Nd is 130Ba16O at mass 146. To control this, 138Ba (the most abundant
Ba isotope) and its oxide at mass 154 was monitored. If the interference is high, the
sample filament is carefully “flashed” at higher temperatures until Ba is burned off. The
presence of La, Pr, Ce does not produce interferences but decreases the ionization
efficiency of Nd. Since the samples studied here had high Nd concentrations this was not
a problem.
Nd measurements were performed on a Finnigan 262 multicollector mass spectrometer.
The measurements were performed in the dynamic mode with one peak jump. The signal
intensities were obtained after online fractionation correction based on the power law and
the 146Nd/144Nd ratio of 0.721903. External reproducibility was measured on the La Jolla
Nd standard. During the period of analysis the average was 143Nd/144Nd = 0.511846+
0.000016 (2 s.d., n = 20).
For the Nd isotope compositions, the initial ε values were calculated using the formula:
⎡ (143Nd / 144Nd ) TS
⎤
− 1⎥ * 1000, were the ratio with the superscript T is the initial
T
143
144
⎣ ( Nd / Nd ) CH
⎦
ratio at the time of eruption, and the subscript CH is the composition of the chondritic
earth at time T.
T
ε Nd
=⎢
Sr and Nd blanks were measured in ICP-MS. An empty beaker was treated like a sample
and processed along with the samples. In the subsequent step, a droplet of 0.25N HCl was
added to the beaker in the loading bench as was done or the samples. This droplet was
108
dried and picked up with 3 ml 1% HNO3 and transferred to a clean vial for the ICP
measurement. At this point all the Sr and Nd present in the beaker representing both the
procedural and reagent blank is in the vial. In order to determine the machine background
the 2% wash solution is used. The blank solution was compared against standard
solutions of 1 and 10 ppt Nd and 100 ppt Sr. The machine background intensity was
around 3% of the Nd intensity in the blank. Thus the blank intensity was sufficiently
higher than the machine background. During the blank measurements it was ensured that
the machine background stays sufficiently low (<10 cps on 146Nd was the primary
monitor). The two blanks determined this way were 9.5 pg and 11.8 pg. This blank
contribution is trivial with respect to the Nd concentrations used here (always several
hundred nanograms). For Sr the two blanks measured were 120 and 130 pg.
A.2. Pb chemical separation and mass spectrometry
For Pb isotopic determination approximately 100 mg sample were dissolved in HF/HNO3
(3:1). Samples were put on a hot plate in a clean air flow box for at least 24h at 1000C to
complete the dissolution. They are then opened and let dry at that temperature
approximately 1000C to evaporate the HF. Lead chemistry was performed following the
techniques described in Abouchami et al., 1999. The samples were then dried three times
in 1 ml of concentrated HNO3, two times in 6N HCl and two times in 1 ml of 9 N HBr.
After dissolution and drying, samples are picked with 0.5 to 1 ml of Solution A (7 ml
H2O + 1 ml 2N HBr + 2 ml 2.5N HNO3) and loaded in 100μl Teflon columns filled with
anion resin. The bulk of the sample is eluted with 0.5 ml of Solution A and 0.2 ml of
Solution B (7.85 ml H2O + 0.15 ml 2N HBr + 2 ml 2.5N HNO3) and Pb is collected with
0.7 ml Solution B. This process is repeated to ensure a pure Pb fraction.
Pb isotope measurements were performed on the Finnigan 262 MS. Samples were picked
up with 0.2N HBr, mixed with silica gel and H3PO4 and loaded on commercial grade
single Re filaments. Since Pb does not have a pair of non-radiogenic isotopes, no online
fractionation correction can be performed. To establish the range of fractionation and
correct for it, several measurements were performed on the NBS 981 Pb standard. To be
able to accurately compare the fractionation observed on the standard to that of the
samples, both standards and samples were run at the same temperatures (1230 to 12500C)
and at the same load (50-75 ng of Pb). The technique involved a fully automated heat up
and run in order to further control fractionation variations arising from different heating
up and running procedures between standards and samples.
The results for the NBS 981 Pd standard were (n = 14):
208
Pb/204Pb = 36.523(+0.023 1 standard deviation)
207
Pb/204Pb = 15.435(+ 0.007)
206
Pb/204Pb = 16.894(+ 0.005)
These values result in a fractionation correction (Todt, et. al., 1993) of 0.12% per amu
determined from analysis of 14 standards.
109
Pb blanks were measured with the ICP-MS. An empty beaker was treated as a sample and
processed along with the samples. At the end, a droplet of 0.2N HBr was added to the
beaker in the loading bench as was done or the samples. This droplet was dried and
picked up with 3 ml 1% HNO3 and transferred to a clean vial for the ICP measurement.
The two measured Pb blanks were 0.3 ng and 0.27 ng.
A.3. ICP-MS analyses
30 mg of sample powder was weighed in a Savillex beakers and dissolved in 3:1
HF:HNO3 mixture for about 48 hours at about 800C. After 2 days the samples are
evaporated to dryness at about 1000C, re-dissolved in 7N HNO3 and left in an ultrasonic
bath for 30 minutes. At this stage most samples were completely dissolved. In case of
incomplete dissolution samples were heated to 175-2000C for several hours and placed in
an ultrasonic bath for ~10 minute. This process was repeated several times. This assured
complete dissolution of all samples. Samples were then transferred to clean HDPE bottles
(pre-cleaned with 1:1 HNO3 for several days), diluted with 2% HNO3 for a total dissolved
solid content of 100 ppm.
Inductively coupled plasma mass spectrometry (ICP-MS) (Finnigan MAT ELEMENT
high-resolution ICP-MS) at the NHMFL was used to determine the concentration of 32
trace elements. CD1E interface (“guard electrode” or platinum shield between the load
coil and torch) was used to attain a sensitivity of about 700,000 cps/ppb 115In for sample
gas and auxiliary gas flow rates of about 1 l/min. A Teflon spray chamber and Teflon
tubing was used along with a Teflon nebulizer from Elemental Scientific, the later was set
at a self aspirating mode with a flow rate of 100 μl/min. The washout characteristics
improved significantly by use of Teflon for the inlet system compared to a conventional
setup consisting of a combination of Tygon tubing and glass nebulizer and spray
chamber. Replacing glass by Teflon allowed the use of a mixture of 2% HNO3-0.03N HF
as cleaning solution in between sample runs. The background for “sticky” elements such
as Nb, Zr, Hf, Pb, Th, and U dropped significantly by a factor of 20-200 after using HF in
the cleaning solution and generally improved the washout characteristics. The blank
levels for most elements were <10-100 cps. A 90-second sample uptake time was chosen
in combination with a washout time of 120 seconds between samples.
Samples were measured in small sequences, consisting of 5 unknowns (4 samples and 1
blank) bracketed by two rock standards, (in this case BIR-1 for mafic rock samples and
G2 for felsic rock samples). Thus the analysis time was kept short and the effects of
signal drift were considerably reduced. Signal drift was usually <5% between the first
and last sample of each sequence. Signal drift was corrected based on an internal
standardization with Indium (all samples and standards are spiked with In to a
concentration of 1 ppb) and a linear interpolation between the external rock standards for
each analyzed isotope. The reference concentrations of BIR-1 are Jochum et al., 2005. To
check for accuracy and precision, repeat measurements were made on one unknown, and
on the rock standard BHVO-I. Based on repeat measurements of the unknown samples
and the rock standard precision for most elements is about 2% (1σ, Table A1).
110
Reproducibility for the rare earth elements (REE) is generally better than 2%. For Nb, Ta,
Zr, Y, Hf, Rb, Sr, and Th, reproducibility is generally better than 4-5%, and Sc, Ti, V, Cr,
Co, Ni and U reproduced to better than 5-7% (Table A1). The average concentration of
BHVO-I agrees well with the values given by Jochum et al. (2005) (Table A1).
Using CD1E interface has both advantages and disadvantages. The CD1E interface
improves sensitivity about five times and thus sample solutions with low total dissolved
solid content can be used which in turn improves the signal drift to a considerable degree
(about four times less signal drift with a 250 ppm solution compared to 500-1000 ppm
solutions over the courses of a two hour measurement). The disadvantage of using CD1E
interface is that it can increase oxide formation up to two times. During the course of
analysis the major oxide interferences on some of the rare earth elements (REE) such as
Eu, Gd and Tb were135BaO on 151Eu, 141PrO on 157Gd, 143NdO on 159Tb. The problem of
oxide formation was dealt in the following ways: a) oxide interferences on samples
enriched in heavy rare earth elements (HREE) are usually insignificant up to oxide
formation levels of several percent, b) oxide formation is approximately constant between
samples with similar matrices. Thus using matrix-matched external standard with similar
Ba, Pr and Nd concentrations compared to the analyzed samples (as in case of BIR-1 and
amphibolites), oxide formation is to a large degree cancelled out by applying the external
rock standard correction. However, especially where samples are more enriched in light
REE and very incompatible elements than the external rock standard, the effects of oxide
formation can be significant, as, for example, between BIR-1 and a LREE enriched rock
such as BHVO-I. Table 1A compares our results for the BHVO-I standard with those
measured by Jochum et al. (2005). Though majority of the elements measured in BHVOI agree to within the analytical uncertainty of the values of Jochum et al. (2005),
measured Gd concentrations in BHVO-I are significantly higher.
A.4. LA-ICP-MS Single Zircon Analysis (method description provided by Dr.
Victor Valencia at University of Arizona, Tuscon)
Samples were collected from GFG, MRG and a thin meta-sandstone unit interlayered
with the PCF amphibolites. At each sample locality, we collected 1–2 kg of fresh whole
rock. Samples were prepared for analysis using standard crushing and separation
techniques, including heavy liquids and magnetic separation. Apparently inclusion-free
zircons were then hand-picked under a binocular microscope. At least 100 zircons from
each sample were mounted in epoxy and polished.
Single zircon crystals were analyzed in polished sections with a Micromass Isoprobe
multicollector ICP-MS equipped with nine Faraday collectors, an axial Daly detector,
and four ion-counting channels (Kidder et al., 2003) at the University of Arizona,
Tucson. The LA-ICP-MS analyses involve ablation of zircon with a New Wave DUV193
Excimer laser (operating at a wavelength of 193 nm) using a spot diameter of 35 to 50
microns. The ablated material is carried in argon gas into the plasma source of a
Micromass Isoprobe, which is equipped with a flight tube of sufficient width that U, Th,
and Pb isotopes are measured simultaneously. All measurements are made in static
mode, using Faraday detectors for 238U, 232Th, 208Pb–206Pb, and an ion-counting channel
111
for 204Pb. Ion yields are 1 mV per ppm. Each analysis consists of one 20-second
integration on peaks with the laser off (for backgrounds), 20 one-second integrations
with the laser firing, and a 30-second delay to purge the previous sample and prepare for
the next analysis. The ablation pit is 20 microns in depth.
Common Pb (initial-age corrected) correction is performed by using the measured 204Pb
and assuming an initial Pb composition from Stacey and Kramers (1975) (with
uncertainties of 1.0 for 206Pb/204Pb and 0.3 for 207Pb/204Pb). Measurement of 204Pb is
unaffected by the presence of 204Hg because backgrounds are measured on peaks
(thereby subtracting any background 204Hg and 204Pb), and because very little Hg is
present in the argon gas.
For each analysis, the errors in determining 206Pb/238U and 206Pb/204Pb result in a
measurement error of several percent (at 2-sigma level) in the 206Pb/238U age. The errors
in measurement of 206Pb/207Pb are substantially larger for younger grains due to low
intensity of the 207Pb signal. The 207Pb/235U and 206Pb/207Pb ages for younger grains
accordingly have large uncertainties. Interelement fractionation of Pb/U is generally
<20%, whereas isotopic fractionation of Pb is generally <5%. In-run analysis of
fragments of a large zircon crystal from a Sri Lanka pegmatite (e.g. Dickinson and
Gehrels, 2003) with known age of 564 ± 4 Ma (2-sigma error) is used to correct for this
fractionation (generally run every third measurement). The uncertainty resulting from the
calibration correction is generally 3% (2-sigma) for both 207Pb/206Pb and 206Pb/238U
ages.
The crystallization ages reported in this thesis are weighted averages of individual spot
analyses. The Paleozoic ages interpreted from the ICP-MS analyses are based on
206
Pb/238U ratios because errors of the 207Pb/235U and 206Pb/207Pb ratios are significantly
greater. This is due primarily to the low intensity (commonly <1 mV) of the 207Pb signal
from these young, U-poor grains. For grains older than 800 Ma, both 207Pb/206Pb and
206
Pb/238U are reported.
A.5. Rb/Sr ratios measured in ICP-MS
All measurements for Galts Ferry Gneiss samples with Rb/Sr ratio less than 5 were
performed on a Finnigan Element 1. Sample preparations were carried out in a Class 300
clean laboratory at the NHMFL. Gravimetric standards were prepared from commercial
Rb and Sr 10.00±0.05 µg/ml stock solutions (High-Purity Standards™). The error
introduced by mixing and dilution of the stock solutions was trivial compared to the
uncertainty on the reported accuracy of the gravimetric solutions (±0.5%).
Rock samples were digested in Savillex™ PFA beakers using PFA-distilled HF-HNO3
acids. The HF was dried down and the residual salts taken up in concentrated HNO3.
These solutions were diluted in quartz-distilled ultrapure water to 2% HNO3 solutions,
stored in acid-washed 125ml LDPE bottles.
Solutions were introduced into the ICP source using an ESI™ low-flow 100 µl/min PFA
nebulizer with an ESI™ PFA spray chamber. A CETAC ASX-100 autosampler was used
112
for automated sample handling. Analyte concentrations were set at levels sufficient to
provide count rates >107 cps that were measured in Analog mode on the Element1
detection system. An analysis of a blank solution (2% HNO3) was performed to correct
the Analog background and to subtract off any memory contributions and 86Kr+
interferences (<104 cps). The magnet was settled at mass 85, and the peaks of interest
(m/e= 85, 86, 87, and 88) were acquired using the EScan option by scanning the
Element’s accelerating voltage in Low Resolution mode (R=300). An isobaric
interference of 87Rb on 87Sr limited the value of that peak. Ratios were calculated for
85
Rb/86Sr, 85Rb/88Sr and 86Sr/88Sr. Measured ion intensity ratios are plotted in Figure A1
against the gravimetric values of the standard solutions. A calibration curve is obtained
that converts measured ion intensity ratios to concentration ratios. Figure A1 shows the
calibration curves obtained for a representative set of measurements of the gravimetric
standards for both 85Rb/86Sr and 85Rb/88Sr. Both ratios provide excellent measures of the
Rb/Sr concentration ratio of the solution, so the final reported Rb/Sr is an average of the
two measurements. To appreciate the precision obtained, Figure A1 also shows the %
difference defined as,
⎡ ( Rb / Sr ) measured
⎤
%difference = ⎢
− 1⎥ * 100,
⎣⎢ ( Rb / Sr ) gravimetric
⎦⎥
as a function of the gravimetric value for Rb/Sr ratios derived from 85Rb/86Sr and
85
Rb/88Sr ratios. Precisions on some average ratios approach ±0.2%, and precision was
always better than ±0.4%. Further, precision of the 86Sr/88Sr was about ±0.4%, as well.
A.6. Rb/Sr ratio measurement by isotope dilution.
The Mulberry Rock Gneiss samples had a huge variation of Rb/Sr ratio and thus standard
isotope dilution technique was used to determine the 87Rb/86Sr ratios. Whole rock
samples were cut into small pieces with a rock saw to remove any traces of weathered
surface. The piece were washed in deionized water in a sonic bath for about 15 minutes,
rinsed in methanol and air dried. The pieces were pulverized in a Siebtecknic rotary
grinder fitted with a tungsten carbide mortar and approximately 100 mg of crushed
samples weighed in teflon screw cap beakers and spiked with a mixed spike of 87Rb - 84Sr
and dissolved in distilled 3:1 HF:HNO3 and similar chemical procedure was followed as
described in A1. The sample solution was then loaded into 12 ml glass column with
Dowex AG50W X-8 (200-400) cation exchange resin and Rb and Sr were eluted in 2.5 N
HCl. Sr measurement was performed on a Finnigan MAT262 Thermal Ionization Mass
Spectrometer following the same procedure described in A1. Rb was measured on the
same instrument between 900 and 1000 degrees centigrade. Three blank analysis of Sr
yielded 0.12 nanograms, 0.13 nanograms, 0.196 nanograms of Sr and two Rb bank
yielded 0.005 microgram and 0.002 microgram of Rb. These values are considered
insignificant for all analysis. In Isoplot 3 (Ludwig, 2001) program used for age
calculation the age data experimental errors of 0.01% for (87Sr/86Sr)N ratios and 1% for
87
Rb/86Sr ratios were used. Age calculation were made using a 87Rb decay constant of
1.42*10-11 year-1.
113
A.7. Rb/Sr ratio measurement by LA-ICP-MS
Feldspar grains with well preserved crystal faces were hand picked under a binocular
microscope. Preferably unaltered orthoclase feldspar grains were picked (identified by
their flesh pink color) so that the major oxide specially the SiO2 concentration can be
estimated for internal standard. The feldspar grains were washed in water and methanol,
mounted in epoxy and left overnight for the epoxy to dry. The epoxy mount was then
polished in 240-321-400-600 grit sand papers and finally in 1 micron diamond polish.
Polished surfaces were analyzed for major and trace elements by laser ablation
inductively coupled plasma mass spectrometry (LA-ICP-MS) at the NHMFL Plasma
Analytical Facility. A New Wave UP213 laser ablation system (213 nm UV) connected
to a Finnigan Element™ ICP mass spectrometer was used for the measurements. The
sample chamber was flushed with helium gas at 800 ml/min, and additional argon makeup gas (980 ml/min) was teed into the sample line to the ICP torch.
The sample surface was scanned in raster mode at a rate of 5 µm/s using a 20 µm
diameter beam, with 10 Hz repetition rate and 50% power output. The mass spectrometer
recorded intensities from the ablated material at the mass values of interest in a series of
275 mass sweeps. Each line scan was 1 mm in length. All of the measured isotopes (7Li,
23
Na, 25Mg, 27Al, 29Si, 39K, 43Ca, 49Ti, 53Cr, 55Mn, 57Fe, 85Rb, 88Sr, 133Cs, 138Ba and 208Pb)
were recorded in the same set of analyses. Blank subtractions (average of 2
measurements with the laser off) were made. Following blank subtraction and correction
by the instrumental sensitivity factors determined on MPI-DING (ATHO-G) reference
glass, the elemental ratios to SiO2 were converted to concentrations assuming SiO2=70%
in orthoclase, then normalized to 100%. Detection limits for each analysis were
calculated from the 3σ uncertainty of blanks run during the analytical session.
Standards used to convert intensities to concentrations included the MPI-DING (ATHOG) reference glass (Jochum et al. 2006). Table A2 gives elemental concentrations for the
standards used in the calibration. Relative sensitivity factors (RSF)
⎡C
⎤
element / C SiO2
⎢
⎥
RSF =
⎢ I element / I Si 29 ⎥
⎣
⎦
for each of the elements analyzed were calculated from the concentrations given in Table
A2 and from measured background-subtracted intensity ratios. Background-subtracted
intensity ratios were then converted to elemental ratios (normalized to SiO2) and
multiplyed by the RSFs. For the Feldspar major elements were calculated from oxide
ratios of Si, Ti, Al, Fe, Mg, Ca, Na, K and Mn normalized to 100%, neglecting minor Cr.
114
Table A1: Repeat analysis of one felsic sample (D2) and USGS standard BHVO1 run as
an unknown is compared with the published concentrations from Jochum et al. (2005).
D2
Li
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Yb
Lu
Hf
Ta
Pb
Th
U
Ti
V
Cr
Co
Ni
Cu
Zn
Ga
D2-rerun
conc (ppm) conc (ppm)
6.79
7.02
127
123
181
182
14.03
14.54
537
564
8.59
8.88
2.94
3.03
478
479
25.98
26.18
53.64
53.96
5.76
5.92
20.04
20.63
3.53
3.55
0.59
0.58
1.48
1.41
0.28
0.30
2.51
2.59
0.50
0.52
1.50
1.55
1.25
1.28
0.23
0.23
16.45
17.02
0.73
0.77
10.04
10.34
11.78
12.12
2.67
2.72
1138
1182
31.70
33.71
130
123
5.87
5.70
28.40
26.77
56.48
53.92
19.42
20.31
3.41
3.55
% difference
3.23%
-3.15%
0.67%
3.56%
4.73%
3.33%
2.95%
0.26%
0.77%
0.59%
2.73%
2.89%
0.51%
-0.68%
-4.88%
5.04%
3.10%
2.73%
3.33%
2.40%
1.67%
3.33%
4.90%
2.84%
2.80%
1.71%
3.70%
5.98%
-5.75%
-3.00%
-6.10%
-4.75%
4.40%
3.95%
115
BHVO1
BHVO1
measured
published
conc (ppm) conc (ppm)
9.02
384
26.34
168
18.06
0.09
129
15.41
37.46
5.27
24.22
5.90
1.96
5.94
0.93
5.33
0.98
2.52
2.06
0.28
4.11
1.11
2.09
1.16
0.40
8.26
395
26.4
174
17.7
0.102
132
15.4
37.9
5.3
24.7
6.13
2.1
6.39
0.959
5.38
0.969
2.54
1.99
0.271
4.26
1.17
2.13
1.23
0.409
% difference
8.42%
-2.97%
-0.23%
-3.54%
2.00%
-8.63%
-2.65%
0.08%
-1.19%
-0.65%
-1.98%
-3.88%
-7.16%
-7.54%
-3.12%
-1.03%
0.66%
-0.90%
3.39%
3.21%
-3.63%
-5.41%
-1.71%
-6.30%
-1.56%
Table A2: Concentrations and relative sensitivity factors (RSFs) for elements analyzed in
this study. Major element oxides in wt. %, trace elements in ppm. Concentrations from
Jochum et al. (2006).
Concentration Sensitivity
Factor
SiO2 (%)
TiO2
Al2O3
FeO
MgO
CaO
K2O
Na2O
MnO
Cr (ppm)
Li
Rb
Sr
Cs
Ba
Pb
75.6
0.255
12.2
3.27
0.103
1.7
2.64
3.75
0.106
6.1
28.6
65.3
94.1
1.08
547
5.67
1
5.31
34.4
2.92
4.09
0.1450
0.0239
131
115
0.0377
0.0101
0.0201
0.0149
0.0295
0.0138
0.0128
Table A3: End member composition of the MORB and crust used for mixing calculation
to produce the Deccan tholeiites.
Charnockites
Basic Granulite Basic Granulite Intermediate Granulite
Weaver et al., 1978 Weaver et al., 1978 Weaver et al., 1978
Rb
Ba
Th
Nb
La
Ce
Pb
Sr
Zr
Ti
Y
100
879
19
18
36
67
11
166
277
2218.15
23
17
400
3
13
28
51
7
210
99
5995
41
N-MORB
Weaver et al., 1978
Hofmann, 1988
4
5
77
3
4
9
19
5
113
56
5575.35
20
320
7
13
39
71
9
180
158
9412.15
30
1.26
13.87
0.19
3.51
3.90
12.00
0.49
113.20
104.24
9681.93
35.82
116
16.00
14.00
y = 7.297245x + 0.011180
R2 = 0.999998
85Rb/86Sr
12.00
10.00
8.00
6.00
4.00
2.00
0.00
0.00
0.50
1.00
1.50
2.00
2.50
Rb/Sr gravim etric
Figure A1. Measured ion
intensity ratios are plotted
against the gravimetric
values of the standard
solutions
(top
two
diagrams).
Percent difference as a
function of the gravimetric
value for Rb/Sr ratios
derived from 85Rb/86Sr and
85
Rb/88Sr ratios (bottom
diagram).
2.00
1.80
y = 0.901111x + 0.000924
R2 = 0.999995
1.60
85Rb/86Sr
1.40
1.20
1.00
0.80
0.60
0.40
0.20
0.00
0.00
0.50
1.00
1.50
2.00
2.50
2.0
2.5
Rb/Sr gravim etric
0.50%
0.40%
% diff 85Rb/86Sr
0.30%
% diff 85Rb/88Sr
% difference
0.20%
0.10%
0.00%
-0.10%
0.0
0.5
1.0
1.5
-0.20%
-0.30%
-0.40%
-0.50%
Rb/Sr grav ime tric
117
SAMPLE LOCATIONS
Sample UTM_X UTM_Y
RA1021 569864 3608902
RA422A 565758 3642576
RA1034 574925 3602673
RA536 548477 3632004
RA401 558126 3643810
RA701B
RA587
RA1005
562161 3645773
565771 3648668
548889 3637505
RA550
557502 3637769
RA1029
RA677
SM 639
570609 3606141
585733 3661572
T18S, R10E, SE1/4,
NE1/4, Sec 11
T20S, R6E, NE1/4
NE1/4 Sec 9
T20S, R6E, NE1/4
NE1/4 Sec 30
T22N, R14E, SE1/4
SW1/4 Sec 11
T20S, R6E, NE1/4
NE1/4 Sec 9
T20S, R6E, NW1/4
NE1/4 Sec21
T23N, R14E, SE1/4
SE1/4 Sec 34
T19S, R7E, SE1/4
SE1/4 Sec 24
T20S, R6E, NW1/4
NW1/4 Sec 28
SM 1100
M 68
M 243
M 268
M 271
M 272
M 201B
M 67
D1
D2
BG 995
BG 999
Y23C
Y24
Y131
NT911
BH166
591872
589674
591868
589664
694526
694602
694507
694405
702698
3676201
3678330
3676204
3678337
3763654
3762308
3763688
3763844
3770971
SC318G
TA7
US41A
Y126
738617
726207
708751
695269
3758025
3733565
3774172
3762310
Quadrangle
Elmore
Flag Mt.
Wetumpka
Mitchell Dam
Mitchell Dam
NW
Flag Mt.
Flag Mt.
Mitchell Dam
NW
Mitchell Dam
NW
Wetumpka
Goodwater
Micaville
Unit
Emuckfaw-Heard
Higgins Ferry
Emuckfaw-Heard
Higgins Ferry
Higgins Ferry
Description
Garnet quartzite
Garnet graphite quartzite
Metagreywacke
Graphitic quartzite
Graphitic quartzite
Higgins Ferry
Higgins Ferry
Higgins Ferry
Metagreywacke
Metagreywacke
Ga+Mu+Bi+Pl schist
Higgins Ferry
Garnet quartzite
Emuckfaw-Heard
Higgins Ferry
Metagreywacke
Graphitic quartzite
Kyanite-garnet schist
Porters Gap
Hillabee Greenstone
Massive greenstone
Porters Gap
Jamison East
Hillabee Greenstone Massive greenstone w/sulfides
and diabasic texture
Hillabee Greenstone
Greenstone/mafic phyllite
Porters Gap
Hillabee Greenstone
Greenstone
Porters Gap
Hillabee Greenstone
Greenstone
Jamison East
Hillabee Greenstone
Foliated greenstone/mafic
phyllite
Foliated greenstone
Lineville West Hillabee Greenstone
Porters Gap
Hillabee Dacite
Metadacite
Bulls Gap
Bulls Gap
Bulls Gap
Bulls Gap
Yorkville
Yorkville
Yorkville
Hillabee Dacite
Hillabee Dacite
Hillabee Dacite
Hillabee Dacite
Galts Ferry Gneiss
Galts Ferry Gneiss
Galts Ferry Gneiss
Galts Ferry Gneiss
Galts Ferry Gneiss
Metadacite
Metadacite
Metadacite
Metadacite
Trondhjemite
Trondhjemite
Trondhjemite
Trondhjemite
Trondhjemite
Galts Ferry Gneiss
Galts Ferry Gneiss
Galts Ferry Gneiss
Pumpkinvine
Amphibolite
Pumpkinvine
Amphibolite
Pumpkinvine
Amphibolite
Pumpkinvine
Amphibolite
Trondhjemite
Trondhjemite
Trondhjemite
Amphibolite
Burnt Hickory
Ridge
South Canton
Taylorsville
Acworth
Yorkville
Y172
693064 3756302
Yorkville
TA16
722287 3734143
Taylorsville
TA17
722104 3739266
Taylorsville
118
Amphibolite
Amphibolite
Amphibolite
Sample UTM_X UTM_Y
Quadrangle
SC254
739757 3758454
South Canton
SC318A
738617 3758025
AC78
709193 3774914
BH57
707261 3773328
SC 25
Sc 33
DR 136
NG 057
MRsamples
685058 3737238
690404 3738984
691697 3746739
692162 3744326
TPGT 1
TPGT 2
TPGT 3
TPGT 4
TPGT
6,7,8,9
TPGT 1115
413293 4056664
413583 4056310
412209 4057522
412512 460952
412512 460952
404378 4052232
Unit
Description
Pumpkinvine
Amphibolite
Amphibolite
South Canton
Pumpkinvine
Amphibolite
Amphibolite
Acworth
Pumpkinvine
Amphibolite
Amphibolite
Burnt Hickory
Pumpkinvine
Amphibolite
Ridge
Amphibolite
Canton Schist
Garnet muscovite schist
Canton Schist
Garnet muscovite schist
Drakestown
Canton Schist
Garnet muscovite schist
New Georgia
Canton Schist
Garnet muscovite schist
New Georgia
Mulberry Rock
Granite samples collceted from
Gneiss
quarry
New Georgia
Mulberry Rock
Granite samples collceted from
Gneiss
quarry
Treas Piedras Tres Piedras Granite
Granite
Treas Piedras Tres Piedras Granite
Granite
Treas Piedras Tres Piedras Granite
Granite
Treas Piedras Tres Piedras Granite
Granite
Treas Piedras Tres Piedras Granite
Granite
Las Tablas
119
Tres Piedras Granite
Granite
100 μ
100 μ
Top: Perthites in Tres Piedras Granite samples.
Bottom: Plagioclase feldspar in Tres Piedras Granite samples.
120
100 μ
Sericitised feldspar in Tres Piedras Granite samples.
121
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BIOGRAPHICAL SKETCH
Reshmi Das joined the Department of Geological Sciences, Florida State University (GS,
FSU) in Fall 2001. Prior to that she completed her Bachelor of Science and Master of
Science programs at the University of Calcutta, India. She has presented papers at various
conferences organized by the Geological Society of America (GSA) and the American
Geophysical Union (AGU). Reshmi received several grants from Educational Mapping
Program (EDMAP), GSA, FSU-Congress of Graduate Students (COGS) etc. Her
research interests include geochemistry, structural geology, and petrology. Reshmi also
served as the instructor for the Dynamic Earth course at GS, FSU for several terms and
was nominated for the FSU Outstanding Teaching Assistant Award 2005.
As a part of her doctoral dissertation Reshmi did three projects on tectonic evolution of
southern Appalachian, isotope homogenization of strontium in northern New Mexico,
and geochemical nature of the first pulses of Deccan lava in India.
Reshmi now joins the National High Magnetic Field Laboratory (NHMFL), Tallahassee,
Florida as a postdoctoral fellow.
Reshmi loves to travel, paint, and read.
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