2013 Sed Geol

This article appeared in a journal published by Elsevier. The attachedcopy is furnished to the author for internal non-commercial researchand education use, including for instruction at the authors institution

and sharing with colleagues.

Other uses, including reproduction and distribution, or selling orlicensing copies, or posting to personal, institutional or third party

websites are prohibited.

In most cases authors are permitted to post their version of thearticle (e.g. in Word or Tex form) to their personal website orinstitutional repository. Authors requiring further information

regarding Elsevier’s archiving and manuscript policies areencouraged to visit:

http://www.elsevier.com/authorsrights

http://www.elsevier.com/authorsrights

Author's personal copy

Early Oligocene paleosols of the Dagshai Formation, India: A record ofthe oldest tropical weathering in the Himalayan foreland

Pankaj Srivastava a,⁎, Subhra Patel a, Nandita Singh a, Toshienla Jamir a, Nandan Kumar a,Manini Aruche a, Ramesh C. Patel b

a Department of Geology, University of Delhi, Delhi-110 007, Indiab Department of Geophysics, Kurukshetra University, Kurukshetra-136119, India

a b s t r a c ta r t i c l e i n f o

Article history:Received 22 February 2013Received in revised form 17 May 2013Accepted 26 May 2013Available online 5 June 2013

Editor: J. Knight

Keywords:Early OligoceneHimalayan ForelandPaleosolsDagshai Formation

This study reports paleopedological features of the fossil soils that formed during the earliest phase of conti-nental sedimentation in the Himalayan foreland. The fluvial sequence of the Dagshai Formation (31.6 ±3.9 Ma to 30.3 ± 3.9 Ma) exposed along the Koshaliya River, NW Himalaya, contains four pedofacies(named Pedofacies A–D) of ferruginous paleosol sequences contained within overbank sediments. TheDagshai Formation unconformably overlies the marine Subathu Formation. Pedofacies A consists of 3–4well-developed ferruginous paleosols overlain by gray sandstone beds. Pedofacies B–D are marked by a pro-gressive decrease in pedogenesis. These paleosols occur as 0.5 m to 1.5 m thick Bw/Bt/Btk/Bk/Bss horizonsthat are marked by extensive development of rhizoliths, pedogenic carbonate, and iron-rich claypedofeatures that correspond to modern Entisols, Inceptisols, Alfisols and Vertisols. Based on early Oligocenepaleogeographic position of the northward-drifting Indian Plate, it is inferred that these paleosols wereformed at ~18°N paleolatitude in the Dagshai sub-basin in the Himalayan foreland. Micromorphology, geo-chemical analyses, weathering indices, and stable isotope composition of paleosols indicate tropical climate(paleoprecipitation of 947–1256 mm and paleotemperature of ~25 °C) with an initial phase of monsoonalconditions during pedogenesis. These paleoclimatic conditions favored C3 paleovegetation immediatelyafter the transition from greenhouse to icehouse conditions.

© 2013 Elsevier B.V. All rights reserved.

1. Introduction

The northward drift of the Indian Plate culminated in collision ofIndia–Asia about ~50 Myr (Klootwijk et al., 1992; Gaina et al., 2007;Kumar et al., 2007; Kent and Muttoni, 2008). This convergence notonly closed the Tethyan Ocean Basin but also resulted in a pronouncedchange frommarine to continental sedimentation all along the Hima-layan foreland (Warwick et al., 1998; Johnson et al., 1999; Welcommeet al., 2001). The dramatic shift frommarine to continental sedimenta-tion is well recorded in the peripheral Himalayan foreland basins, suchas the Subathu–Dagshai basin. The “passage beds” consisting of varie-gated shale of the Subathu Formation andwhite-gray sandstone of theDagshai Formation demarcate the transition frommarine to continen-tal conditions (Mathur, 1977; Fig. 1). The fluvial deposits of theDagshai Formation consist of sandstone–mudstone sequences withpaleosols (Kumar et al., 2008; Fig. 2).

The transitional sedimentary record of the Subathu and Dagshaiformations has been intensively studied with regard to tectonics,paleoclimate, sedimentology, geochronology and paleobiology

(Raiverman and Raman, 1971; Bhatia, 1982; Sahni et al., 1983;Bhatia, 2000; Singh, 2003; Najman et al., 2004; Bhatia and Bhargava,2006; Najman, 2006; Bera et al., 2008; Kumar et al., 2008; Bera etal., 2010a,b; Ravikant et al., 2011). Despite the significant informationknown about the marine to continental transition, paleopedologicalconditions of these strata remain a major gap in our understandingof the oldest continental sedimentary record from the Himalayanforeland. This article presents a macromorphological, micromorpho-logical, and geochemical study of the ferruginous paleosols from theDagshai Formation.

The entire succession of the Dagshai Formation studied herein isexposed along the Koshaliya River on Kalka-Shimla Highway, NWHimalaya (Fig. 1), and has been explored for detailed characterizationof the fossil soils. The paleosols from 18 litho-sections were studiedfor macroscopic and micromorphological details for their degree of de-velopment (maturity). Bulk geochemical analyses of the paleosols alongwith stable carbon and oxygen isotopic values of pedogenic carbonatewere carried out to support themacro- andmicromorphological obser-vations, calculate weathering indices, and infer paleoenvironmentalconditions. These paleopedogenic data are used to infer the responseof early Oligocene weathering and pedogenic development in theDagshai basin to climate changes induced by northward drift of theIndian Plate.

Sedimentary Geology 294 (2013) 142–156

⁎ Corresponding author.E-mail address: [email protected] (P. Srivastava).

0037-0738/$ – see front matter © 2013 Elsevier B.V. All rights reserved.http://dx.doi.org/10.1016/j.sedgeo.2013.05.011

Contents lists available at ScienceDirect

Sedimentary Geology

j ourna l homepage: www.e lsev ie r .com/ locate /sedgeo


2. Geological setting

The Himalayan mountain range is the most prominent and activeintracontinental range in the world. Continued convergence of the

Indian Plate has resulted in progressive development of a series ofmajor thrust/faults (Gansser, 1964; Molnar and Tapponnier, 1975;Wobus et al., 2005). The outer part of this range in the south ismarked by an active foreland basin that resulted from thrust loading

Fig. 1. (a) Location of India, (b) location of study area (Subathu basin) in the Himalayan foreland (after Gansser, 1964). (c) General geological map of the Subathu–Dagshai Forma-tions of the Himalayan foreland (after Raiverman, 1979; Raiverman et al., 1983), (d) location of the measured litho-sections (D1 to D18) from Dagshai Formation along theKoshaliya River, NW Himalaya.

143P. Srivastava et al. / Sedimentary Geology 294 (2013) 142–156


and subsidence with synorogenic sedimentation from the hinterland(Raiverman et al., 1983; Burbank et al., 1996). The Himalayan Fore-land Basin is bounded by the Main Boundary Thrust to the northand the Indo-Gangetic Plains in the south (Parkash et al., 1980;Raiverman et al., 1983; Burbank et al., 1996). The terrestrial sedimen-tary record preserved in this foreland is linked with exhumation his-tory of the hinterland, southward propagation of boundary thrusts,and progressive shift of sediment provenance (Ravikant et al.,2011). The synorogenic sediments in the marine Subathu sub-basinrepresent the basal part of the foreland succession and were sourcedfrommixed Tethyan and Lesser Himalayan rocks to the north and alsofrom Cratonic rocks of the Indian shield from the south (Ravikantet al., 2011). The overlying continental Dagshai Formation consistsof red mudstone and white/gray sandstone deposited by meanderingrivers and records drastic changes in depositional environment andsediment provenance (Bhatia and Bhargava, 2006; Najman, 2006;Bera et al., 2008; Ravikant et al., 2011).

The Subathu sub-basin (Fig. 1) successions in theHimalayan forelandrepresent the last phase of marine sedimentation during Paleocene–Eocene (DeCelles and Giles, 1996; Najman, 2006; Bera et al., 2008,2010a; Ravikant et al., 2011). The marine sediments are overlain by flu-vial sediments of the Dagshai and Kasauli Formations, with a possiblehiatus of long duration (Najman et al., 1993, 1994, 2004; Bhatia andBhargava, 2006; Najman, 2007). Based on the occurrence of large ben-thic foraminifera, the sediments near the Subathu–Dagshai boundarywere assigned an age of 44 Ma with bio-chronological continuity acrossthe boundary (Mathur, 1978; Bhatia and Bhargava, 2006). The deposi-tional age of the lowermost and uppermost parts of the Dagshai forma-tion has been dated to 31.6 ± 3.9 Ma and 30.3 ± 3.9 Ma, respectively,using detrital-zircon fission-track (ZFT) analysis (Jain et al., 2009). Theage of ~31 Ma for strata in the basal part of the Dagshai Formation im-plies a major hiatus between marine (Subathu Fm.) and continental(Dagshai Fm.) sedimentation (Najman et al., 2004; Najman, 2006; Jainet al., 2009), which is supported by detailed sedimentary facies and pa-leomagnetic studies (Sangode et al., 2005; Kumar et al., 2008).

It is important here to consider the influence of post-pedogenic dia-genesis with regard to the paleoenvironmental potential of the DagshaiFormation sediments. It is estimated that maximum burial of theDagshai Formation occurred during the deposition of the Kasauli For-mation and Siwalik Group (Neogene), and later by thrusting of LesserHimalayan rocks along the Main Boundary Thrust (Karunakaran andRanga Rao, 1976). Overall a minimum of 7–7.5 km depth of burial hasbeen estimated based on thickness of overlying sedimentary forma-tions. This corresponds to aminimum of ~140 °C temperature of burial,considering a geothermal gradient of ~20 °C km−1 for the Siwalik basin(Mugnier et al., 1995). The estimated burial temperature based onvitrinite reflectance and spore color indices indicates a range of 140–195 °C (Najman et al., 2004).

3. Material and methods

In the present study, 18 litho-sections were measured to recordmacromorphological features of the paleosols (Soil Survey Staff, 1999)over the entire fluvial succession of the Dagshai Formation exposedalong the Koshaliya river (Fig. 1). The 65 paleosols from the 18litho-sections were examined for color, structure, texture, Fe/Mn fea-tures, CaCO3 and rhizoliths to characterize them as master horizons(A, B, C, etc.) following Keys to Soil Taxonomy (Soil Survey Staff, 1999,2006). The paleosols of 50–100 cm thickness were sub-sampled at20–30 cm intervals for detailed analyses.

For micromorphological study, 65 thin-sections (4 × 6 cm) from20 paleosols (Fig. 2) across the 18 litho-sections were made followingMurphy (1986) and described under the petrographic microscope fol-lowing the terminology given by Bullock et al. (1985). The degree ofpedogenic development was assessed on the basis of maturity ofpaleopedogenic features following criteria established for macro-and micromorphological features (Bullock et al., 1985; Birkeland,1999). Accordingly, well-developed, moderately-developed, andweakly developed paleosols are marked by strongly to poorlyexpressed pedofeatures, respectively, similar to those in modern

Fig. 2. Stratigraphic section measured along Koshaliya River (after Kumar et al., 2008) and correlating paleosols (D1 to D18). Details of the four-pedofacies (A–D) in terms of degreeof soil development (poorly, moderately, and well-developed) and bounding sediments (gray sandstone). The log of the 18 measured sections (D1–D18), this also includes detailsof the representative thin sections and the key-pedofeatures shown in Fig. 3(a–d). Age of Dagshai Formation after Jain et al. (2009).

144 P. Srivastava et al. / Sedimentary Geology 294 (2013) 142–156


polygenetic soils and late Quaternary paleosols of the Gangetic Plains(Srivastava et al., 1994; Kumar et al., 1996; Srivastava and Parkash,2002; Srivastava et al., 2010). We have used a pedofacies approachto organize paleosols of Dagshai Formation. This approach is akin tothat used for pedofacies described from overbank deposits in whichpaleosols experience lateral changes in properties as a result of dis-tance (during formation) from areas of relatively high sediment accu-mulation (Bown and Kraus, 1987).

Forty samples of Dagshai Formation paleosols representing a rangeof well-developed to poorly developed examples were analyzed forbulk geochemistry. For this, bulk samples of the paleosols were groundto about ~200 μmmeshusing Tungsten Carbide balls in a TemaMill andthen oven-dried at 110 °C. Pressed powder pellets were prepared bymixing 7–8 g of the sample with 4–5 drops of polyvinyl alcohol as abinding agent (Saini et al., 2000, 2002). The prepared pellets werekept in a hot boxmaintained at 60 °C to drive off themoisture. All sam-ples including the reference standards used for calibration were madeunder identical conditions to avoid any error introduced by samplepreparation. Chemical analyses were performed on a WavelengthDispersive X-ray Fluorescence Spectrometer (Siemens SRS-3000) atWadia Institute of Himalayan Geology, Dehradun, India. Calibration ofthe instrument was done using a set of well characterized internationalreference samples (SO-1: Soil-CCRMP-Canada; GSS-1 and GSS-4: Soil-IGGE-China, GXR-2, GXR-6: Soil-USGS-USA; SCO-1, SGR-1, SDO-1:Shale-USGS-USA; MAG-1: Marine mud-USGS-USA; GSD-9, GSD-10:Sediment-IGGE-China; GSR-6: Limestone-IGGE-China; BCS-267: Silicabrick: BAS-UK). The precision and accuracy of instrumental perfor-mance were checked using international reference samples SO-1 andSDO-1. The accuracy of measurements was found to be better than2–5% in case of major oxides and 12% for most of the trace elements.Precision in terms of standard deviation of repeated measurements isb3%. In order to assess chemical weathering, chemical index of alter-ation without potash [CIA-K = 100 ∗ Al2O3/(Al2O3 + CaO + Na2O)]was determined for paleosol horizons (Bt/Bw) showing well-developed pedofeatures (Maynard, 1992). An empirical relationship be-tween CIA-K and annual precipitation (P in mm/yr = 221e0.0197 ∗ CIA-K)based on hydrolysis of silicates during weathering causing enrichmentof Al at the expense of Ca, Mg, Na and K in modern soils (R2 = 0.72;S.E. = ±182 mm) proposed by Sheldon et al. (2002) has been usedto estimate mean annual precipitation (MAP) for Dagshai paleosols.For paleosols with vertic features (Bss horizons) a refinement overCIA-K as CALMAG = Al2O3/(Al2O3 + CaO + MgO) ∗ 100 proposedby Nordt and Driese (2010) was used for estimation of MAP (P inmm/yr = 22.69 ∗ CALMAG − 435.8; R2 = 0.9; S.E. = ±108 mm).

The stable isotope composition of 11 pedogenic calcium carbonatesamples representing the lowermost part of the Dagshai Formation

was also analyzed. A micro-drill system was used to collect the sam-ples from carbonate nodules of the paleosols. The samples were pul-verized to b100 μm size and reacted with orthophosphoric acid at70 °C for 2 h in glass vials for characterization of δ13C and δ18O. Theevolved CO2 was measured chromatographically on an Isoprime Con-tinuous Flow Mass Spectrometer at Nuclear Hydrology Laboratory ofNational Institute of Hydrology, Roorkee, India. The stable isotopecomposition is presented in the usual δ-notation as the per mill (‰)deviation of the sample relative to PDB standard with analytical pre-cision of ±0.1‰ for both δ13C and δ18O. The δ13C compositionrecorded for Dagshai paleosols is compared to pedogenic carbonatesforming soils with dominance of C3/C4 vegetation (Koch, 1998;Cerling and Harris, 1999; Breecker et al., 2009). The δ13C values ofpedogenic carbonates from Dagshai paleosols has been used to esti-mate atmospheric pCO2 using the isotope mass balance relationship(Cerling, 1999; Ekart et al., 1999) defined as (pCO2)(atm) =S(z)(δ13Cs − 1.0044δ13Cr − 4.4) / (δ13Ca − δ13Cs). For this estimatethe variables were taken as: S(z) soil respired CO2 = 3382 ppmv and4189 ppmv (after Retallack, 2009: soil respired CO2 = 66D + 588;where D represents uncompacted depth of Bk horizon), carbon isotopecomposition of soil CO2 as δ13Cs = −19.9‰ and 20.2‰, soil respiredCO2 as δ13 Cr = −22.0‰, and atmospheric CO2 as δ13Ca = −3.0.‰,and the coefficient 1.0044 and the constant 4.4 derived from the differ-ence in diffusivity between 13CO2 and 12CO2 (Retallack, 2009).

4. Results

4.1. Macroscopic features of the paleosols

Macroscopic evidence of pedogenic development in fluvial se-quences of the Dagshai Formation includes root traces, blocky struc-ture, color mottling, pedogenic carbonate (PC) nodules, slickensides,and illuvial clay coatings. On the basis of degree of development(Table 1), Dagshai Formation paleosols that are marked by well-developed pedofeatures are compared to QGH4–QGH5 soils of theGangetic Plains formed over 8000–10,000 years. The Dagshai Forma-tion is characterized by an upward decrease in the abundance ofwell-developed paleosols (Figs. 2b, 3). The eighteen litho-sections ex-amined are organized into four pedofacies (Pedofacies-A toPedofacies-D) that occur sequentially over the entire DagshaiFormation (Fig. 2).

Pedofacies-A is marked by 3–4 well-developed paleosols and fewpoorly-developed paleosols (e.g., litho-sections D1–D5, D7). Pedofacies-A mainly occurs in the lower part of the Dagshai Formation as ~4 mthick paleosols in overbank mudstones that are overlain and underlainby 1–2 m thick cross-bedded sandstones (Figs. 2, 3a, 4a). Well-

Table 1Criteria for determining the degree of development in paleosols (after Srivastava et al., 1994, 2010).

Degree of development Criteria Pedofacies in Dagshai Fm. Comparable soils/paleosols in GangeticPlains (after Srivastava et al., 1994, 2010)

Strongly developed >1 m thick B horizon (Bw/Bt/Bss/Bk),well-developed structural elements (peds),common illuvial features (>1%), commonrhizoliths, mottles and pedogenic carbonate

Pedofacies-A(e.g. D1, D2, D3, D4, D5, D7) mainlyin the lowermost part of Dagshai Fm.

QGH4–5 soils with strongly developedpedofeatures formed in 10–8 ka time

Moderately-developed 50–70 cm thick B horizon (Bw/Bss/Bk),moderately-developed structural elements(peds), illuvial features (b1%), commonrhizoliths, mottles and pedogeniccarbonate

Pedofacies-B(e.g. D6, D10, D12, D13, D15) mainlyin middle part of Dagshai Fm.

QGH3 soils with moderately-developedpedofeatures formed in 2.5–5 ka time

Weakly/poorly developed 20–30 cm thick B horizon (Bw/BC/Bk), poorly-developedstructural elements (peds), few rhizoliths,mottles and pedogenic carbonate

Pedofacies-C(e.g. D8, D9, D14) mainly in middle-upper part of Dagshai Fm.

QGH2 soils with weakly developedpedofeatures (0.5–2.5 ka)

Sediments with little orno pedogenic activity

Dominance of sedimentary featureswithout any prominent B horizon, butoccasionally few rhizoliths and mottles

Pedofacies-D(e.g. D11, D14, D18) mainly inuppermost part of Dagshai Fm.

QGH1 soils with little or no pedogenicactivity (b0.5 ka)



developed paleosols within this pedofacies are marked by dusky-redcolor (10R), silty texture, blocky and platy structures with extensiverhizoliths and pedogenic carbonate features (Fig. 4b; Supplementarydata Table A). The root traces occur as elongate, cylindrical featureswhich taper and branch out downwards (Fig. 4). These root traces are15–20 mmwide in upper horizons and 4–5 mmwide in lower horizonsand extend from 0.5 m to 1 m on average and in some cases up to 2 m indepth. These rhizoliths are tubular in form with gray core and red orreddish-brown rims. In addition, many of these rhizoliths are alsomarked by calcareous accumulations as infillings along the root haloes(Fig. 4j). Angular blocky to sub-angular blocky structures are commonbut some of the horizons are marked by slickensides with wedge shapedand platy peds (Fig. 4d, g). The lowermost part of Pedofacies-A is marked

by an erosional surface showing extensive pedorelicts in contact withsandstone with no evidence of pedogenic alteration (Fig. 4b). Thesepedorelicts with fabrics similar to the ferruginous paleosols of the areaoccur as discrete features in transitional horizons showing erosional con-tact with underlying sediments.

Pedofacies-B is characterized by 2–3 well-developed paleosols and3–4 poorly-developed paleosols (e.g., litho-sections D6, D10, D12–13,and D15). Pedofacies-B occurs mainly in the middle part of the DagshaiFormation and preserves fewer (2–3) well-developed paleosols com-pared to pedofacies-A (Figs. 2a, b, 3b).

Further upwards only moderately-developed paleosols occurand are defined as Pedofacies-C (e.g., litho-sections D8–9, D14;Figs. 2, 3). Pedofacies-D defines the uppermost part of the Formation,

Fig. 3. A summarized account of the key -pedofeatures of the paleosols representing Pedofacies-A (a: D2, and D5 paleosols), Pedofacies-B (b: D10, and D12 paleosols), Pedofacies-C(c: D8, and D9 paleosols), and Pedofacies-D (d: D11, and D17 paleosols).



Fig. 4. Macro- and micromorphic details of Pedofacies-A (D2 section paleosols). (a) Outcrop details of well-developed paleosol (marked by arrow) from the lowermost part of theDagshai Formation. (b) Details of the well-developed paleosol (D2, Pedofacies-A) showing macroscopic features (color, rhizocretions, structures etc.) with characteristic horizons(Bw, Bt, Btk, Bss, and BC) bounded by gray sandstone, (c) a part of this paleosol showing rhizoliths as branching of the root trace (marked by arrow), (d) strongly developed angularblocky peds with iron (goethite) mottles (marked by arrow), D2/3, Btk1 horizon, 90 cm depth, (e) thick microlaminated iron-rich clay pedofeatures, D2/3, Btk1 horizon, 90 cmdepth, (f) iron-rich clay coatings along the voids (marked by arrow), D2/2b, Bt horizon, 70 cm depth, (g) strongly developed angular blocky peds from a vertic horizon showingstriations due to pedogenic slickenside, D2/6, Bss horizon 180 cm depth, (h) Micromorphological details of wedge shaped peds from Bss horizon with strongly oriented b-fabricand stress argillans, D2/7, Bssk1 horizon, 200 cm depth, (i) Strongly developed slickenside (ss) marked by parallel striations, wedge shaped peds, and stress clay coatings (markedby arrow) along with pedogenic calcium carbonate (pc), D2/7, Bssk1 horizon, 200 cm depth, (j) A ped from vertic horizon showing a branching rhizocretion filled with dense cal-cium carbonate (arrow), D2/8, Bssk2 horizon, 240 cm, (k) cross section of a root trace showing walls coated with organic matter and iron rich clay along the root channel, note thePart of root showing elongation along the slickenside, D2/8, Bssk2 horizon, 240 cm, (l) pedogenic calcium carbonate (pc) nodule showing dense impure micrite with diffuse bound-aries and fragmentation, D2/4, Btk2 horizon, 110 cm depth, (m) biogenic reworking showing bow shaped fabric of pedogenic calcium carbonate, organic matter, and iron-rich clay,D2/3, Btk1 horizon, 90 cm.



which is dominated by sediments with little or no pedogenic activityand only a few weakly to very weakly developed paleosols (e.g.,litho-sections D11, D14, and D18; Figs. 2, 3, 6).

4.2. Micromorphological features of the paleosols

Representative thin-sections of the four pedofacies are shown inFigs. 4–6. Key-micromorphological pedofeatures such as pedality,b-fabric, bioturbation, iron-rich clay coatings and pedogenic CaCO3

are well preserved in these paleosols and are briefly described below.Pedofacies-A: Micromorphology of the paleosols from litho-section

D2 indicates moderately to strongly developed pedofeatures in 3–4distinct paleosol units separated by erosional surfaces with pedorelicts(Fig. 3a; Supplementary data Table A). A sharp contact with lithic sand-stone defines the upper part of this paleosol sequence (Fig. 4b). The

underlying paleosol (1.1 m thick) shows strong pedality and silty tex-ture with c/f ratio (coarse vs fine ratio at 20 μm limit, defined byBullock et al., 1985) varying from 40/60 to 30/70 (Fig. 4). The coarsefraction consists of silt- and fine sand-sized quartz, mica and lithic frag-ments that are coated with iron-rich clay (Fig. 4e, f). Thick iron-rich clayand fine silty-clay pedofeatures occur as coatings along the voids, rootchannels, and as intercalations (Fig. 4e, f). The birefringence fabric(b-fabric) is defined by well-developed cross and parallel striations(Fig. 4i, k, l). Rhizoliths occurring as pedotubles are marked by the fol-lowing features: (i) a central core infilled by groundmass andpedorelicts, (ii) mineralization by clay minerals, calcium carbonate,and iron oxide, (iii) rarely preserved coalified organic residue, (iv) irondepletion zones around the central core as gray mottles, and (v) iron-oxide rich silty clay/clay coatings (Figs. 4j, 5b, c, k, l). Pedogenic carbon-ate varies from 2 to 5% in the upper horizons to 60–80% in the lower

Fig. 5.Macro- and micromorphological features of well-developed paleosols from Pedofacies-B (D10, D12, D15 sections). (a) A ped showing moderate–strong angular blocky struc-ture with depletion along root trace and iron mottles, D10/4, Bt3 horizon, 90 cm depth, (b) iron-rich clay intercalations along a root channel (arrow), D10/4, Bt horizon, 90 cmdepth, (c) fragmented intercalations of iron-rich clay (arrow), D10/4, Bt3 horizon, 90 cm depth, (d) contrasting fabric due to slickenside showing strong alignment of parallel stri-ated b-fabric and pedogenic calcium carbonate (arrow pointing slickenside), D12/2, Bss horizon, 60 cm depth, (e) moderately dense nodule of pedogenic calcium carbonate (PC)with diffused boundary in the groundmass, D12/3, Bss horizon, 90 cm depth, (f) pedogenic calcium carbonate nodules showing coarsening of the fabric, D12/3, Bss horizon, 90 cmdepth, (g) A ped showing angular blocky structure along with pedogenic calcium carbonate (arrow) and root traces, D15/1, Bw horizon, 30 cm depth, (h) pedogenic calcium car-bonate (pc) with diffuse boundaries in the groundmass, D15/1, Bw horizon, 30 cm depth, (i) mottles of iron oxide showing initial stage of concretion formation with concentricfabric (arrow), D15/1, Bw horizon, 30 cm depth, (j) Moderately–strongly oriented parallel striated b-fabric and clay coatings along planar void, Bss horizon (arrow), D15/2, Bsshorizon, 60 cm depth (k) iron-rich clay coating along the walls of a root channel, D15/3, Bss horizon, 90 cm depth, (l) part of the root channel wall showing fragmentation ofthe coating, D15/3, Bss horizon, 90 cm depth.



horizons (Fig. 4i, l). It is marked by (i) dense micritic nodules with dif-fuse boundaries (Fig. 4i, l), (ii) loose to dense micritic cement in theroot channels, (iii) soft and loose micritic nodules in the groundmasswith floating grain fabric, and (iv) pedorelicts of calcareous mud withbiogenic features and spherulitic calcite. The strongly developedpedofeatures correspond to Alfisols with Bt and Btk horizons. Belowthis, Alfisol vertic features predominate for about 1.2–1.5 m depth, de-fined by strong pedality with angular blocky and wedge shaped peds,slickenside surfaces, stress argillans, parallel striated b-fabric, and elon-gated nodules of pedogenic carbonates (Fig. 4g–i). The upper and lowerparts of this vertic horizon are marked by extensive pedorelicts, whichindicate presence of erosional surfaces, and is underlain by a paleosolshowing features similar to those in the upper part, and correspondsto Inceptisol-Alfisol-like characters. The lowermost part of this paleosolsequence is in contact with underlying gray sandstone and is marked byan erosional surface with relicts of sediments and paleosols. Micromor-phology of the paleosols from litho-section D5 also shows features sim-ilar to those described for D2 (Fig. 3a).

Pedofacies-B:Micromorphology of paleosols from sections D10, D12and D15 shows moderately–strongly developed pedofeatures (Supple-mentary Data Table A). Paleosols from section D10 are marked bystrong pedality, speckled to striated b-fabric, pedorelicts, iron-richsilty-clay coatings along root channels and intercalations (Fig. 5), and10–15% pedogenic CaCO3 in lower parts. The pedofeatures from sectionD12 indicate vertic horizons marked by wedge shaped peds, parallel

striations, slickensides and stress clay coatings along planar voids(Fig. 5d–f). The pedofeatures from section D15 are similar to D10 andcorrespond to Alfisol-Inceptisol (Fig. 5g–l).

Pedofacies-C: Micromorphology of paleosols from section D9 indi-cates weak–moderate development. Thin sections from D9 show weakpedality, sandy groundmass, stipple-speckled b-fabric, rhizocretions,few iron-rich clay intercalations, andpedogenic carbonate that increasesin lower horizons (Supplementary data Table A).

The thin sections from Pedofacies-D are characterized by sporadicoccurrence of pedogenic features (Fig. 6; Supplementary dataTable A). The paleosols in Pedofacies-C and Pedofacies-D correspondto Inceptisol-Entisol like features.

4.3. Bulk rock geochemistry of the paleosols

Bulk rock geochemistry of 10 paleosols from sections D2, D5,D9 and D10 was used to evaluate weathering indices andpaleoenvironmental conditions during pedogenesis. Parent materialfor these paleosols can be defined as C and BC horizons (mudstoneswith little or no pedogenic activity). Major and minor element com-position data are presented as weight percentages of oxides andtrace elements as ppm (see Supplementary data Tables B, D, and E).A cross plot for major oxides vs horizons presented in Fig. 7 providesan assessment of weathering in Dagshai paleosols. The SiO2 and Al2O3

values show a relative decrease in Bt horizons when compared to A,

Fig. 6. Macro- and micromorphic details of Pedofacies-D (D16, D17 paleosols). (a) Outcrop details of the poorly-developed paleosol from Pedofacies-D representing the uppermostpart of the Dagshai Formation, D 16 section, (b) details of the D 16 section showing weakly developed pedogenic features in mudstone (root traces marked by arrow), (c) lithicsandstone showing framework grains (Quartz, Mica, and Rock fragments) with iron and calcareous cements, D16/3, C horizon, 90 cm depth, (d) laminated mudstone with veryweakly developed pedogenic features marked by a few mottles (arrow), D17/5, BC horizon, 160 cm depth, (e) weakly developed pedofeatures marked by iron oxide mottles(arrow) and cross-striated b-fabric, D11/3, Bw2 horizon, 90 cm depth, (f) iron-rich clay intercalation along a root channel (arrow), D11/2, Bw1 horizon, 60 cm depth.



but little or no change when compared to C/BC horizons (Fig. 7).However, it showed a large spread for Fe2O3 values in all horizonswith no distinct trend (Fig. 7). On the other hand horizon distribution

of CaO, MgO and Na2O showed a relative increase in Bt when com-pared to A or C/BC horizons (Fig. 7). This trend is less pronouncedfor Na2O with large spread in different horizons (Fig. 7). The ratio of

Fig. 7. Geochemical characteristics of Dagshai Paleosols showed as variation of SiO2, Al2O3, Fe2O3, CaO, Na2O, MgO, vs different horizons (A/Bss/Bt/Bw/BC/C) of the paleosols. Ratiosfor Al2O3/SiO2 and U/Th and Sr distribution in different horizons (A/Bss/Bt/Bw/BC/C) given in plots show weathering characteristics of the paleosols.



Al2O3/SiO2 values show an increase in Bt horizon when comparedwith other horizons (Fig. 7). Trace element behavior for Sr and theratio of U/Th showed a relative increase in Bt horizons when com-pared with A horizons (Fig. 7). The chemical index of alteration with-out potash (CIA-K) show high values ranging from 77 to 89 for Bt/Bwhorizons that correspond to MAP from 947 ± 182 mm to 1287 ±182 mm (Supplementary data Table B). The CALMAG values deter-mined for the paleosols with vertic features (Bss horizons) rangedfrom 65 to 68, that correspond to MAP from 1039 ± 108 mm to1116 ± 182 mm (Supplementary data Table B).

4.4. Stable isotope geochemistry of the pedogenic carbonates

The stable isotope composition data (δ13C and δ18O) of the D2paleosol (Pedofacies-A) in the lowermost part of the Dagshai forma-tion are plotted in Fig. 8. Both isotopic values are marked by depletionin heavier isotopes (Supplementary data Table C; Fig. 8a). The depthdistribution of δ13C and δ18O values shows three distinct zones:(i) 50–150 cm slight increase, (ii) 150–250 cm slight decrease,and (iii) 250–320 cm slight increase of heavier isotopes correspond-ing to three distinct paleosol units (Alfisols and Vertisols) with ero-sional surfaces (Figs. 3a, 8b). The δ13C composition for Dagshai

paleosols (−9.4‰ to −11.3‰) is compared to soil carbonates withdominance of C3 vegetation (δ13C values:−10‰ to−13‰). Applica-tion of isotope mass balance relationship applied at two Bk horizonsprovides an estimate of atmospheric pCO2 (433 ppmv and633 ppmv) during Dagshai time.

5. Discussion

5.1. Interpretation of the paleopedofeatures of Dagshai Formation paleosols

The macromorphological and micromorphological features de-scribed for the four pedofacies reflect well-developed ferruginouspaleosols in lower parts of the formation and weakly developedpaleosols in upper parts (Figs. 2–6). Well-developed paleosols showingBt, Btk, Bss, Bssk horizons are marked by strong pedality, cross and par-allel striated b-fabric, bioturbation, mineral weathering, rhizocretions,ferruginous nodules/concretions, and pedogenic carbonates (e.g., D2and D5 in Pedofacies-A; D10 in Pedofacies-B). The pedofeatures occur-ring as thick-ferruginous silty-clay coatings and as intercalations ofsilty-clay (e.g., Figs. 4e, f, k, 5j, k) indicate substantial illuviation and for-mation of Bt horizons in these paleosols in wet conditions (Chandranet al., 2005; Srivastava et al., 2009).

The pedogenic calcium carbonate (PC) occurs in greater amounts(25–30% reaching up to 60–80%) in lower horizons compared toupper horizons (nil to 5%) of calcareous profiles (e.g., D2). This trendis more pronounced in Pedofacies-A compared with Pedofacies-B, C,andD that aremarkedby anoverall decrease in PC. The pedogenic fabricof these carbonate features correspond to Stage II-III (Wieder andYaalon, 1982; Machette, 1985) showing impure micritic nodules withdiffused boundaries (Figs. 4l, 5h). The PC described here is similar toLate Quaternary paleosols and Holocene soils of the Himalayan fore-land, and their formation has been ascribed to lowering of pCO2 byevapotranspiration in semiarid to subhumid climates (Srivastava,2001; Srivastava et al., 2007, 2010). Presence of PC alongwith clay coat-ings suggests that PC formation occurred in subhumid to humid climaticconditions. Dominance of PC in lower horizons (>50 cm) with dissolu-tion features in these paleosols is similar to PC inmodern Vertisols fromsubhumid to humid regions (Srivastava et al., 2002). Presence of spher-ulitic calcite indicates the role ofmicroorganisms (bacterial action) dur-ing precipitation of calcium carbonate that is similar to pedogeniccarbonates forming in modern tropical soils (Braissant et al., 2003).Pedogenic slickensides (Bss horizons) observed in these paleosols sug-gest shrinking and expansion of swelling-clays due to wetting and dry-ing (Wilding and Tessier, 1988; Pal et al., 2001; Srivastava et al., 2002).Presence of smectite-rich clays reported from the Dagshai Formation(Raiverman and Suresh, 1997; Raiverman, 2002) suggests that slicken-sides in these paleosols occurred due to shrinking-swelling of smectite.The morphological expression of pedogenic slickensides is smooth,polished and striated surfaces (cf., Figs. 4g, h, i, 5j). Low-relief (50–100 cm) variations in paleosol surfaces in the field suggest hummockymicrotopography with predominance of vertic features (cf., D2paleosol). This may be expressed as low-relief pseudo-anticlines andpseudo-synclines due to the presence of swelling-clays (Joeckel, 1989;Caudill et al., 1996; Driese et al., 2003; Hillier et al., 2007).

Abundance of rhizoliths in these paleosols can be ascribed to theirformation after woody vegetation prevalent during early Oligocenepedogenesis. The rhizoliths are comparable to ichnofossils fromother terrestrial sediments such as elongate gray mottles with redrims occurring in (well-drained) ferruginous paleosols from BighornBasin, USA (Kraus and Hasiotis, 2006).

In the Dagshai Formation, red colored mottles and concretions areindicative of hematite (Fe2O3) and goethite (α-FeOOH) in the paleosols(Figs. 4c, d, 5i). In view of deep burial (~7 km) it is expected thatremobilization and precipitation of iron-oxide were possibly caused bydiagenetic alteration, but this could have been limited because structuralelements, textural pedofeatures, mottling, pedogenic calcium carbonate

Fig. 8. Stable isotope (δ13C and δ18O) composition of the paleosols from the lower mostpart of the Dagshai Formation (D2 Paleosols, D2/2-D2/11 horizons) representingwell-developed pedofeatures from Pedofacies A, (a) Cross plot of the δ13C and δ18Ovalues for the paleosols from D2 section, (b) δ13C and δ18O values versus depth varia-tion for the well-developed paleosols showing three distinct zones corresponding toAlfisols and Vertisols.



and swell-shrink features are still well-preserved and the Dagshai sedi-ments have a significant proportion of smectite (Raiverman and Suresh,1997; Raiverman, 2002) and unaltered lithic fragments of very low-grade metamorphism (Chaudhri, 1975; Najman and Garzanti, 2000).The XRD studies confirmed the dominance of hematite and goethite inthese paleosols. This supports limited influence of diagenesis, whichotherwise could have affected all forms of iron oxides in a similar wayrather than producing a mix of hematite and goethite (PiPujol andBuurman, 1994). A pedogenic origin of the red color in these ferruginouspaleosols is in agreementwith the soils that develop in tropical and sub-tropical environments (Kampf and Schwertmann, 1982), attributed towell-drained and oxidizing conditions (Macedo and Bryant, 1987;PiPujol and Buurman, 1994; Kraus and Hasiotis, 2006).

Morphological expression of Dagshai paleosols shows stacking ofsoil types on top of each other due to compound or cumulic pedogen-esis (Kraus and Aslan, 1993). The paleosols in lower parts representcompound paleosols which typically show overprinting and weldednature of well-developed paleosols (Fig. 3). However, in upper partssimple paleosols with weakly developed pedogenic features domi-nate that are possibly caused by varying pace of sedimentation andpedogenesis (Srivastava et al., 2010). The upward decrease of well-developed and iron-rich clayey paleosols suggests an overall increaseof sedimentation in comparison with pedogenesis, which is in agree-ment with a study describing the sedimentation of the Dagshai For-mation by flooding of interfluves in coastal areas (Kumar et al.,2008). The dominant white/gray sandstone in the lowermost DagshaiFormation has been interpreted as evidence of accelerated uplift andrapid sea-level fall during the transition from marine (Subathu) tocontinental (Dagshai) sedimentation (Bera et al., 2008, 2010a). Theoverlying paleosols represents the oldest terrestrial record from theNW Himalayas (Kumar et al., 2008) and mark a critical period ofEarth history (~31 Ma) following the Eocene-Oligocene transition(Fig. 9) and the change from greenhouse to icehouse conditions(Kent and Muttoni, 2008; Zachos et al., 2008).

5.2. Bulk rock geochemistry and weathering characters of the paleosols

Whole rock analysis of paleosols can be useful to evaluate pedo-genic processes and weathering indices (Nesbitt and Young, 1982;Maynard, 1992; Retallack, 1997a, 1997b, 1997c; Sheldon et al., 2002;Driese et al., 2005; Kraus and Riggins, 2007; Sheldon and Tabor,2009; Nordt and Driese, 2010; Adams et al., 2011; Retallack andDilcher, 2012). The ratios of the major oxides are used to quantifythe intensity of chemical weathering (Nesbitt and Young, 1982;Maynard, 1992; Sheldon et al., 2002; Nordt and Driese, 2010). Thechemical index of alteration (CIA) proposed by Nesbitt and Young(1982) measures the degree of rock weathering and can assessweathering and pedogenic development in silicate rocks (Sheldonand Tabor, 2009). However, in view of possible addition of potassiumdue to burial diagenesis, this index was modified (CIA-K) and K wasomitted from the CIA (Maynard, 1992). The CIA-K expression hasbeen widely used to estimate the mean annual precipitation (MAP)and the degree of chemical weathering in paleosols without vertic fea-tures (Sheldon et al., 2002; Driese et al., 2003; Nordt et al., 2006;Prochnow et al., 2006). But for the paleosols with vertic features (Bsshorizons), a refinement over CIA-K as CALMAG weathering index de-fined as Al2O3 / (Al2O3 + CaO + MgO) ∗ 100 has been found to im-prove rainfall estimation (Nordt and Driese, 2010).

The MAP estimate based on CIA-K values (for Bt/Bw horizons) andCALMAG values (for Bss horizons) ranges from 947 mm to 1287 mm,suggesting wet conditions during pedogenesis. Presence of pedogeniccarbonate (PC) in these paleosols with wet conditions is in accordwith PC of Vertisols from humid regions at >50 cm depth with disso-lution features (Srivastava et al., 2002). Micromorphology of thepaleosols showed PCwith dissolution features in lower horizons (Sup-plementary Data Table A). The geochemical composition of thepaleosols in Bt/Bw horizons when comparedwith C/BC and A horizonssuggests moderate-high weathering due to leaching of bases and con-centration of Al, Ti and Fe (Supplementary data Table E; Fig. 7). The

Fig. 9. (a) A correlation of the fluvial sedimentary record of the Himalayan foreland (Dagshai Formation) with marine oxygen isotope record using magnetic polarity scale (afterCande and Kent, 1992, 1995). Age of the paleosols is based on Zircon Fission Track (ZFT) data from basal and upper part of the Dagshai Formation (Jain et al., 2009). The marineisotopic data are from benthic foraminifera (Miller et al., 1987). The Oi-1 and Oi-2 represent Eocene–Oligocene cooling events (Miller et al., 1991). The well-developed ferruginouspaleosols of the Dagshai Formation at ~31 Ma represent early Oligocene warm-humid conditions, (b) Paleolatitudinal reconstruction of the Indian Plate including the location ofDagshai Formation at 50 Ma, 30 Ma, and Present (after Besse and Courtillot, 2002), suggesting tropical climatic conditions during early Oligocene times (~31 Ma), when paleosolsformed.



paleosols with argillic horizons show more leaching and strongerweathering in comparison with paleosols with vertic horizons. Theratio of Al2O3/SiO2 shows higher values for Bt/Btk horizons in compar-isonwith BC/Bwhorizons suggesting formation of pedogenic clay dur-ing weathering (Fig. 7).

Trace element abundance can vary considerably during pedogenesisand has been used to assess weathering conditions (Retallack, 2001a,2001b; Kahmann et al., 2008; Sheldon et al., 2009). Themost commonlyused trace elements are Sr, U, Th, Nb and Zr (Sheldon and Tabor, 2009).In general Sr is soluble and leaches readily from soils. In the presentstudy Sr concentrations show an increase with depth (Supplementarydata Table E; Fig. 7), suggesting that chemical weathering was intense.U and Th are both relatively immobile, but under strong chemicalweathering U is leached from the profile while Th remains (Sheldonand Tabor, 2009). We record higher U/Th ratios in Bt/Btk than Bw/BChorizons (Supplementary data Table E; Fig. 7) which suggest that Uleached from upper horizons and was redistributed to lower horizons.

5.3. Stable isotope composition of pedogenic carbonates and paleoclimate

5.3.1. Paleovegetation record in Dagshai paleosolsThe stable isotope composition (δ13C, δ18O) of PC is potentially use-

ful to track vegetation and climatic conditions during paleosol forma-tion (Pendall et al., 1994; Quade and Cerling, 1995; Cerling et al.,1997; Srivastava, 2001; Behrensmeyer et al., 2007). The soil CO2 pro-vides a major control on precipitation of PC, which in turn is basicallycontrolled by the proportion of the biomass (C3 and C4 plants) in theecosystem (Cerling and Quade, 1993). Pedogenic carbonates formingin soils dominated by C3 vegetationwill bemarked by δ13C values rang-ing from −10‰ to −13‰ whereas soils dominated by C4 vegetationwill form soil-carbonate with δ13C values ranging from +1‰ to +4‰(Koch, 1998; Cerling and Harris, 1999; Breecker et al., 2009). In thepresent study, δ13C values of PC range from −9.4‰ to −11.3‰ andare thus closely related to C3 vegetation (Fig. 8a). This is also supportedby macro and micromorphological features that indicate dominance ofwoodland vegetation with thick roots (~20–30 mm wide) extendingover 2 m depth. This is in agreement with the paleobotanical evidenceof fossil leaf impressions of Ficus spp. in fluvial sequences of the DagshaiFormation, which is an evergreen to occasionally deciduous, moderateto large tree distributed over tropical humid areas (Mathur et al.,1996; Srivastava et al., 2011). The paleolatitudinal position (~18°N) ofthe Dagshai sub-basin during the early Oligocene confirms tropical con-ditions (Fig. 9), which are also seen in paleovegetation from fluvial sed-iments of Oligocene age from NorthWest and North-East Himalaya,with their paleolatitudinal position at 10–15°N in comparison withtheir present position at 27–29°N (Metais et al., 2009; Martin et al.,2011; Srivastava et al., 2012). The carbon isotopic composition of thepaleosol carbonates is in close agreement with δ13C values (−10.1 to−10.3‰) from fossil enamel teeth of Oligocene (28–33 Ma) age fromthe North-West Himalaya (Martin et al., 2011), and is consistent withpaleosol carbonates from early Oligocene paleosols that indicate C3type vegetation at this time (Cerling, 1992; Bestland et al., 1997; Alroyet al., 2000; Retallack et al., 2000; Zachos et al., 2001; Koch et al., 2003).

5.3.2. Atmospheric pCO2 during the early OligoceneUnderstanding the relationship of atmospheric pCO2 and climate

provides crucial information for climate change over geological time(Breecker et al., 2009, 2010; Cotton and Sheldon, 2012). Paleoclimateduring the Cenozoic is marked by a declining trend of atmosphericpCO2 with values of 1000 to 1500 ppmv during the Eocene that de-creased to modern levels in the Oligocene (Berner and Kothavala,2001; Zachos et al., 2001; DeConto and Pollard, 2003; Pagani et al.,2005). In the present study an estimate of pCO2 of 433 ppmv and630 ppmv is much lower than the global estimate of ~1200 ppmv dur-ing the early Oligocene (Ekart et al., 1999). This could be explained con-sidering the fact that soil-respired CO2 [S(z)]was previously assumed to

bemuch higher while calculating paleo-atmospheric pCO2 (Breecker etal., 2009, 2010; Cotton and Sheldon, 2012). This commonly led to anoverestimation of paleo-atmospheric pCO2 (Breecker et al., 2009,2010; Cotton and Sheldon, 2012). Soil-respired CO2 S(z) is directly pro-portional to the predicted atmospheric value, so assuming an incorrectS(z) can lead to large errors in estimation of atmospheric pCO2 (Cottonand Sheldon, 2012). For example, Ekart et al. (1999) assumed S(z) valueof 5000 ppmv for the Phaneorozoic, based on S(z) estimates of modernsoils for mean growing seasons (Brook et al., 1983; Solomon andCerling, 1987). However, a recent study shows that pedogenic carbon-ate growth is seasonal and occurs in warm, dry periods whensoil-respired CO2 [S(z)] ismuch lower than in themean growing season(Breecker et al., 2009). In the present study, our estimate of atmosphericpCO2 is based on S(z) values for the warm, dry season (Retallack, 2009),and is in good agreement with estimates based on stable isotope com-position of alkenones (Pagani et al., 2005) and from Ginkgo stomatalindex (Retallack, 2009).

5.3.3. Diagenetic influence and δ18O values of pedogenic carbonateIn view of burial depth (~7 km) of the Dagshai paleosols, it is im-

portant to consider the influence of diagenesis on isotopic composi-tion of the pedogenic carbonates (Cerling, 1991; Mora et al., 1996).The diagenetic alteration makes it difficult to use δ18O values fromthe present study alone unless there is a record of unaltered δ18Ovalues for these paleosols. In light of this observation, we refer tothe isotopic composition of tooth enamel from large mammals (pro-boscideans, rhinocerotoids and equids), which is more stable thanpedogenic carbonate and likely to preserve original physicochemicalconditions even after deep burial (Lee-Thorp and van der Merwe,1987; Cerling and Harris, 1999; Passey et al., 2005; Martin et al.,2011). The record of the isotopic composition of teeth enamel(rhinocertoids) from fluvial sediments of Bugti Hills (NW Himalaya)that formed during the Dagshai interval shows δ13C values from−11.9 to −9.5‰ (mean −10.5‰) and δ18O values from −3.2 to−2.2‰ (mean −2.5‰) (Martin et al., 2011). This shows that diage-netic alteration of pedogenic carbonates caused a significant shift ofthe oxygen isotope composition but no change in carbon isotopecomposition. The diagenetic influence on stable isotope compositionhas been explored in several studies which demonstrate that burialdiagenesis can cause additional depletion of δ18O values up to 10‰in pedogenic carbonates (Cerling, 1991; Quade et al., 1995). Consider-ing this effect, we can make an assessment of unaltered δ18O valuesfor paleosol carbonates, which are in the range from −7.9 to−2.7‰ (mean −5.1‰). This is in agreement with δ18O values forthe dense micrite representing unaltered pristine carbonates fromthe Dagshai basin (Bera et al., 2010b). These recalculated values re-flect the isotopic composition of paleoprecipitation, because the oxy-gen isotopic composition of soil carbonates is primarily controlled bymeteoric water (Cerling, 1984; Cerling and Quade, 1993; Rowley etal., 2001). This corresponds to the δ18O values of present-day tropics(values of −4.0 to −7.0‰; IAEA, 1992; Hoffmann and Heimann,1997) suggesting a tropical paleoclimate.

Paleotemperature during formation of the pedogenic carbonates canbe estimated using the δ18O-temperature relationship curve (Cerling,1984). Based on this,mean annual temperature of ~20.5 °C has been es-timated. We have also used the soil-temperature scale based on the re-lationship of oxygen isotope fractionation between calcite and water atlow temperatures to estimate paleotemperature (Friedman and O’Neil,1977; Kim and O’Neil, 1997). Using this relationshipwith δ18O values of−6.0‰ reported for precipitation at New Delhi (IAEA, 1992) and thelowest δ18O value (−7.9‰) of pedogenic carbonatewith a fractionationfactor (α = 1.0284), yields a paleotemperature estimate of 24.5 °C,which is higher than mean annual temperature. The period of paleosolformation in the Dagshai sub-basin around the Eocene-Oligocene tran-sition is marked by lowered global temperature (Zachos et al., 2001).The latitudinal climatic gradient suggests temperature anomalies to be



highest (20 °C to 40 °C) at high (>60°N) latitudes, moderate (10 °C to20 °C) inmid-latitudes, andwith little or no variation in low latitudes atthis time (Utescher et al., 2011). This suggests little or no variation ofpaleotemperature in the tropics. Thus, formation of the paleosols mayhave occurred with paleotemperature close to modern temperaturevalues.

5.4. Incipient monsoonal conditions during early Oligocene

In order to track the pattern of paleoprecipitationwe refer to the for-mation of pedogenic carbonate in sub-soils. Pedogenic carbonate forma-tion occurs in warm, dry conditions (Birkeland, 1999; Breecker et al.,2009, 2010; Cotton and Sheldon, 2012), and can form during dry sea-sons under monsoon climates. Pedogenic carbonates in modern soilsunder monsoonal conditions are typically marked by a larger spreadcompared to thickness of Bk horizon and ferric-calcite concretions(Sehgal and Stoops, 1972; Retallack, 1991, 2005). Pedogenic carbonatein Dagshai paleosols shows similarity with pedogenic carbonates ofGangetic Plains in the following respects: (i) dense impuremicritic nod-ules, (ii) dissolution and re-precipitation features in Bk horizon, and (iii)spread of pedogenic carbonate beyond the Bk horizons (Srivastava,2001). Based on these observations it is interpreted that formation ofpedogenic carbonates in Dagshai paleosols possibly occurred during ini-tial stages of monsoonal conditions.

This is in contrast with most of the proposed ages for initial mon-soon development in this region, which cluster at 22–23 Ma (Guo etal., 2002; Clift et al., 2008). Although there is no consensus, an oldermonsoonal phase cannot be ruled out completely in light of followingobservations: (i) Shrinking Paratethys and the rising Himalaya in re-sponse to northwardmotion of the Indian Plate influenced atmosphericcirculation over a large scale (Ramstein et al., 1997), (ii) Climate-orogeny studies suggest moderate relief mountains (~2000 m) existedduring the early Oligocene that attained higher elevations (~4500 m)during the late Miocene (Copeland et al., 1987; Richter et al., 1991;Harrison et al., 1992; Ramstein et al., 1997; Dupont-Nivet et al., 2007),(iii) Rise of theHimalaya changed the climatemore towards continentalconditions with increased seasonality (Ramstein et al., 1997), and(iv) During the Oligocene, the southern part of India was locatedbelow the inter-tropical convergence zone (ITCZ) and received highrainfall that was annually distributed, whereas middle and northernIndia experienced 50–60% monsoonal precipitation (Ramstein et al.,1997). These observations suggest that monsoonal conditions duringformation of Dagshai paleosols were weaker than today.

6. Conclusions

Results of this study show the pedological and geochemical char-acters of fossil soils from the oldest terrestrial sediments of the NWHimalayas. The fluvial successions of the Dagshai Formation are char-acterized by four pedofacies that are marked by an upward decreasein well-developed paleosols. Well-developed paleosols are thick(1–1.5 m) showing Bt/Btk/Bss/Bk horizons with extensive rhizoliths,iron-rich clay coatings, and pedogenic carbonates.

Paleolatitudinal reconstruction suggests that the earliest terrestrialsedimentation occurred close to ~18°N during the early Oligocene.The geochemical characters of these paleosols show moderate–strongweathering with leaching of bases and formation of pedogenic clay.The carbon and oxygen isotopic composition of pedogenic calcium car-bonate indicate the prevalence of C3 type vegetation and wet-warm(MAP: 947–1256 mm; MAT: 25 °C) tropical conditions. The paleosolsare marked by pedogenic carbonates that developed under a weakmonsoon but annually distributed high rainfall.

Supplementary data to this article can be found online at http://dx.doi.org/10.1016/j.sedgeo.2013.05.011.

Acknowledgments

We acknowledge financial support by the Department of Scienceand Technology SR/S4/ES-081/2003 to PS. We thank Dr. D. K. Paland Dr. R. Kumar for helpful discussions to write this manuscript.We also thank Dr. B. Kumar and Dr. S. Kumar for the help during sta-ble isotopic studies at National Institute of Hydrology, Roorkee andDr. P. Mukherjee, WIHG, Dehradun for geochemical analysis. Wethank Dr. Steven Driese, an anonymous reviewer, and editor Dr. JasperKnight for thoughtful reviews of the manuscript.

References

Adams, J.A., Kraus, M.J., Wing, S.L., 2011. Evaluating the use of weathering indices fordetermining mean annual precipitation in the ancient stratigraphic record.Palaeogeography, Palaeoclimatology, Palaeoecology 309, 358–366.

Alroy, J.P., Koch, P.L., Zachos, J.C., 2000. Global climate change and North Americanmammalian evolution. In: Erwin, D.H., Wing, S.L. (Eds.), Deep time: Paleobiology'sperspective: Paleobiology, 26, pp. 259–288.

Behrensmeyer, A.K., Quade, J., Cerling, T.E., Kapplelman, J., Khan, I.A., Copleland, P., Roe,L., Hicks, J., Stubble, P., Willis, B.J., Latorre, C., 2007. The structure and rate of lateMiocene expansion of C4 plants; evidence from lateral variation in stable isotopesin paleosols of the Siwalik Group, northern Pakistan. Geological Society of AmericaBulletin 119, 1486–1505.

Bera, M.K., Sarkar, A., Chakraborty, P.P., Loyal, R.S., Sanyal, P., 2008. Marine to continen-tal transition in Himalayan foreland. Geological Society of America Bulletin 120,1214–1232.

Bera, M.K., Sarkar, A., Chakraborty, P.P., Ravikant, V., Choudhary, A.K., 2010a. Forcedregressive shoreface sandstone from Himalayan foreland: implications to earlyHimalayan tectonic evolution. Sedimentary Geology 229, 268–281.

Bera, M.K., Sarkar, A., Tandon, S.K., Samantha, A., Sanyal, P., 2010b. Does burial diagen-esis reset pristine isotopic compositions in paleosol carbonates? Earth and Plane-tary Science Letters 300, 85–100.

Berner, R.A., Kothavala, Z., 2001. GEOCARB III: a revised model of atmospheric CO2 overPhaneorozoic time. American Journal of Science 301, 182–204.

Besse, J., Courtillot, V., 2002. Apparent and true polar wander and the geometry of thegeomagnetic field over the last 200 Myr. Journal of Geophysical Research 107(B11), 2300. http://dx.doi.org/10.1029/2000JB000050.

Bestland, E.A., Retallack, G.J., Swisher III, C.C., 1997. Stepwise climate change recordedin Eocene–Oligocene paleosol sequences from central Oregon. Journal of Geology105, 153–172.

Bhatia, S.B., 1982. Facies, fauna and flora in the lower Tertairy formations of NorthwesternHimalayas: a synthesis. Palaeontological Society of India, Special Publication 1, 8–20.

Bhatia, S.B., 2000. Faunal and floral diversity in the Subathu–Dagshai passage beds: areview. Himalayan Geology 21, 87–98.

Bhatia, S.B., Bhargava, O.N., 2006. Biochronological continuity of the Paleogene sedi-ments in the Himalayan foreland Basin: Palaeontological and other evidences.Journal of Asian Earth Sciences 26, 477–487.

Birkeland, P.W., 1999. Soil and Geomorphology. Oxford University Press (430 pp.).Bown, T., Kraus, M.J., 1987. Integration of channel and floodplain suites, I. developmental

sequence and lateral relations of alluvial paleosols. Journal of Sedimentary Petrology57, 587–601.

Braissant, O., Calilleau, G., Dupraz, C., Verrcchia, E.P., 2003. Bacterially inducedmineraliza-tion of calcium carbonate in terrestrial environments: the role of Exopolysaccharidsand amino acids. Journal of Sedimentary Research 73, 485–490.

Breecker, D.O., Sharp, Z.D., McFadden, L.D., 2009. Seasonal bias in the formation andstable isotopic composition of pedogenic carbonate in modern soils from centralNew Mexico, USA. Geological Society of America Bulletin 121, 630–640.

Breecker, D.O., Sharp, Z.D., McFadden, L.D., 2010. Atmospheric concentrations duringancient greenhouse climates were similar to those predicted for A.D. 2100. Pro-ceedings of the National Academy of Sciences of the United States of America107, 576–580.

Brook, G.A., Folkoff, M.E., Box, E.O., 1983. A world model for soil carbon dioxide. EarthSurface Processes and Landforms 8, 79–88.

Bullock, P., Fedoroff, N., Jongeroius, A., Stoops, G., Tursina, T., 1985. Handbook of SoilThin Section Description. Waine Research Publication, Wolverhampton (152 pp.).

Burbank, D.W., Beck, R.A., Mulder, T., 1996. The Himalayan foreland basin. In: Yin, A.,Harrison, M. (Eds.), The Tectonic Evolution of Asia. Cambridge University Press,pp. 149–188.

Cande, S.C., Kent, D.V., 1992. A new geomagnetic polarity time scale for the late Creta-ceous and Cenozoic. Journal of Geophysical Research 97, 13917–13951.

Cande, S.C., Kent, D.V., 1995. Revised calibration of the geomagnetic polarity time scale forthe late Cretaceous and Cenozoic. Journal of Geophysical Research 100, 6093–6095.

Caudill, M.R., Driese, S.G., Mora, C.I., 1996. Preservation of a paleo-Vertisol and an esti-mate of late Mississippian paleoprecipitation. Journal of Sedimentary Research 66,58–70.

Cerling, T.E., 1984. Stable isotope composition of modern soil carbonate and its rela-tionship to climate. Earth and Planetary Science Letters 71, 229–240.

Cerling, T.E., 1991. Carbon dioxide in atmosphere: evidence from Cenozoic and Meso-zoic paleosols. American Journal of Science 291, 377–400.

Cerling, T.E., 1992. Development of grasslands and savannas in East Africa during theNeogene. Palaeogeography, Palaeoclimatology, Palaeoecology 97, 241–247.



Cerling, T.E., 1999. Stable carbon isotopes in paleosol carbonates. In: Thiry, M., Simm-Coincon, R. (Eds.), Palaeoweathering, Palaeosurfaces and related continentaldeposits. Special Publication of the International Association of Sedimentologists.Blackwells, pp. 43–60.

Cerling, T.E., Harris, J.M., 1999. Carbon isotope fractionation between diet and bioapatitein ungulate mammals and implications for ecological and paleoecological studies.Oecologia 120, 347–363.

Cerling, T.E., Quade, J., 1993. Stable carbon and oxygen isotope in soil carbonates. In:Swart, P.K., Lohmann, K.C., McKenzie, J.A., Savin, S. (Eds.), Climate change in conti-nental isotopic records: American Geophysical Union, pp. 217–231.

Cerling, T.E., Harris, J.M., McFadden, B.J., Leaky, M.G., Quade, J., Eisenmann, J.,Ehleringer, J.R., 1997. Global vegetation change through the Miocene and Pliocene.Nature 389, 153–158.

Chandran, P., Ray, S.K., Bhattacharyya, T., Srivastava, P., Krishanan, P., Pal, D.K., 2005.Lateritic soils of Kerala, India: their mineralogy, genesis, and taxonomy. AustralianJournal of Soil Research 43, 839–852.

Chaudhri, R.S., 1975. Sedimentlogy and genesis of the Cenozoic sediments of North-western Himalayas (India). Geologische Rundschau 64, 958–977.

Clift, P.D., Hodges, K.V., Hannigan, R., Long, H.V., Calves, G., 2008. Correlation of Himalayanexhumation rates and Asian monsoon intensity. Nature Geoscience 1, 875–880.

Copeland, P., Harrison, T.M., Kidd, W.S.F., Ronghua, X., Yuquan, Z., 1987. Rapid earlyMiocene acceleration of uplift in the Gangdese belt, Xizang (southern Tibet), andits bearing on accommodation mechanism of the Indian–Asia collision. Earth andPlanetary Science Letters 86, 240–252.

Cotton, J.M., Sheldon, N.D., 2012. New constraints on using paleosols to reconstructatmospheric pCO2. Geological Society of America Bulletin 124, 1411–1423.

DeCelles, P.G., Giles, K.A., 1996. Foreland basin systems. Basin Research 8, 105–123.DeConto, R.M., Pollard, D., 2003. A coupled climate-ice sheet modeling approach

to the early Cenozoic history of the Antarctic ice sheet. Palaeogeography,Palaeoclimatology, Palaeoecology 198, 39–52.

Driese, S.G., Jacobs, J.R., Nordt, L.C., 2003. Comparison of modern and ancient Vertisolsdeveloped on limestone in terms of their geochemistry and parent material. Sedi-mentary Geology 157, 49–69.

Driese, S.G., Nordt, L.C., Lynn, W.C., Stiles, C.A., Mora, C.I., Wilding, L.P., 2005.Distinguishing climate in the soil record using chemical trends in a Vertisolclimosequence from the Texas Coast Prairie, and application to interpreting Paleo-zoic paleosols in the Appalachian Basin, U.S.A. Journal of Sedimentary Research 75,339–349.

Dupont-Nivet, G., Krijsman, W., Langereis, C.G., Abels, H.A., Dai, S., Fang, X., 2007. Tibet-an plateau aridification linked to global cooling at the Eocene–Oligocene transition.Nature 445, 635–638.

Ekart, A., Cerling, T.E., Montanez, I.P., Tabor, N.J., 1999. A 400 million year carbon isotoperecord of pedogenic carbonate: implications for paleoatmospheric carbon dioxide.American Journal of Science 299, 805–827.

Friedman, I., O’Neil, J.R., 1977. Compilation of stable isotope fractionation factors ofgeochemical interest, In: Fleischer, M. (Ed.), Data of Geochemistry, 6th edition.U.S. Geological Survey Professional Paper, 440-KK, pp. 1–12.

Gaina, C., Muller, R.D., Brown, B., Ishihara, T., 2007. Breakup and early seafloor spread-ing between India and Antarctica. Geophysical Journal International 170, 151–169.

Gansser, A., 1964. Geology of the Himalayas. Interscience, London (289 pp.).Guo, Z.T., Ruddiman, W.F., Hao, W.F., Wu, Q.Z., Qiao, H.B., Zhu, Y.S., Peng, R.X., Wei, S.Z.,

Yaun, B.S., Liu, T.S., 2002. Onset of Asian desertification by 22 Myr ago inferredfrom loess deposits in China. Nature 416, 156–163.

Harrison, T.M., Copeland, P., Kid, W.S.F., Yin, A., 1992. Raising Tibet. Science 255,1663–1670.

Hillier, R.D., Marriott, S.B., Williams, B.P.J., Wright, V.P., 2007. Possible climate variabilityin the Lower Old Red Sandstone, Coingar Pit Sandstone Member (early Devonian),South Wales, U.K. Sedimentary Geology 202, 35–57.

Hoffmann, G., Heimann, M., 1997. Water isotope modeling in the Asian Monsoon region.Quaternary International 37, 115–128.

IAEA, 1992. Statistical treatment of data on environmental isotopes in precipitation.IAEA, Vienna 781 (STI/DOC/10/331).

Jain, A.K., Lal, N., Sulemani, B., Awasthi, A.K., Singh, S., Kumar, R., Kumar, D., 2009. Detrital-zircon fission-track ages from the Lower Cenozoic sediments, NW Himalayan fore-land basin: clues for exhumation and denudation of the Himalaya during theIndia–Asia collision. Geological Society of America Bulletin 121, 519–535.

Joeckel, R.M., 1989. Geomorphology of a Pennsylvanian land surface: pedogenesis inthe Rock Lake Shale Member, southeastern Nebraska. Journal of SedimentaryPetrology 59, 469–481.

Johnson, E.A., Warwick, P.D., Roberts, S.B., Khan, I.H., 1999. Lithofacies, depositional envi-ronments and regional stratigraphy of Lower Eocene Ghazij Formation, Balochistan,Pakistan. United States Geological Survey Professional Paper 1599, 1–76.

Kahmann, J.A., Seaman III, J., Driese, S.G., 2008. Evaluating trace elements as paleoclimateindicators: multivariate statistical analysis of the late Mississippian PenningtonFormation Paleosols, Kentucky, U.S.A. Journal of Geology 116, 254–268.

Kampf, N., Schwertmann, U., 1982. Goethite and hematite in a climosequence in south-ern Brazil and their application in classification of kaolinitic soils. Geoderma 29,27–39.

Karunakaran, C., Ranga Rao, A., 1976. Status of exploration for hydrocarbons in theHimalayan region, Contributions to stratigraphy and structure: Himalayan GeologySeminar, Section III, Oil and Natural Gas Resources. Geological Survey of India,Miscellaneous Publications 41, 1–66.

Kent, D.V., Muttoni, G., 2008. Equatorial convergence of India and early Cenozoic cli-mate trends. Proceedings of the National Academy of Sciences 105, 16065–16070.

Kim, S.T., O'Neil, J.R., 1997. Equilibrium and nonequilibrium oxygen isotope effects insynthetic carbonates. Geochimica et Cosmochimica Acta 61, 3461–3475.

Klootwijk, C.T., Gee, J.S., Peirce, J.W., Smith, G.M., McFadden, P.L., 1992. An early India–Asia contact: paleomagnetic constraints from Ninety east Ridge, ODP leg 121.Geology 20, 395–398.

Koch, P.L., 1998. Isotopic reconstruction of past continental environments. AnnualReview of Earth and Planetary Sciences 26, 573–612.

Koch, P.L., Clyde, W.C., Hepple, R.P., Fogel, M.L., Wing, S.L., Zachos, J.C., 2003. Carbonand Oxygen isotope records from paleosols spanning the Paleocene-Eoceneboundary, Bighorn Basin, Wyoming. In: Wing, S.L., Gingerich, P.D., Schimitz, B.,Thomas, E. (Eds.), Causes and consequences of globally warm climates in theearly Paleogene: Geological Society of America Special Paper, 369, pp. 49–64.

Kraus, M.J., Aslan, A., 1993. Eocene hydromorphic paleosols: significance for interpretingancient floodplain processes. Journal of Sedimentary Petrology 63, 453–463.

Kraus, M.J., Hasiotis, S.T., 2006. Significance of different modes of rhizoliths preserva-tion to interpreting paleoenvironmental and paleohydrologic settings: examplesfrom Paleogene paleosols, Bighorn Basin, Wyoming, U.S.A. Journal of SedimentaryResearch 76, 633–646.

Kraus, M.J., Riggins, S., 2007. Transient drying during the Paleocene–Eocene ThermalMaximum (PETM): analysis of paleosols in the Bighorn Basin, Wyoming.Palaeogeography, Palaeoclimatology, Palaeoecology 245, 444–461.

Kumar, S., Parkash, B., Manchanda, M.L., Singhvi, A.K., Srivastava, P., 1996. Holocenelandform and soil evolution of the western Gangetic Plains: implications ofneotectonics and climate. Zeitschrift für Geomorphologie N.F. 103, 283–312.

Kumar, P., Yuan, X., Kumar, M.R., Kind, R., Li, X., Chadha, R.K., 2007. The rapid drift ofthe Indian tectonic plate. Nature 449, 894–897.

Kumar, R., Sangode, S.J., Ghosh, S.K., Sinha, S., 2008. Marine to fluvial transition anderosional hiatus in the Paleogene sediments of NW Himalayan Foreland Basin,India. Himalayan Geology 29, 147–160.

Lee-Thorp, J.A., van der Merwe, N.J., 1987. Carbon isotope analysis for fossil boneapatite. South African Journal of Science 83, 712–715.

Macedo, J., Bryant, R.B., 1987. Preferential microbial reduction of hematite over goethitein a Brazilian Oxisol. Soil Science Society of America Journal 53, 1114–1118.

Machette, M.N., 1985. Calcic soils of the southwestern United States. In: Wiede, D.L.(Ed.), Soils and Quaternary geology of the southwestern United States: GeologicalSociety of America Special Paper, 203, pp. 1–22.

Martin, C., Betalleb, I., Antoine, P.O., 2011. Pakistan mammal tooth stable isotopes showpaleoclimatic and paleoenvironmental changes since the early Oligocene.Palaeogeography, Palaeoclimatology, Palaeoecology 311, 19–29.

Mathur, N.S., 1977. Age of the Tal and Subathu formations in the Garhwal region, UttarPradesh, India. Bulletin of the Indian Geologists Association 10, 21–27.

Mathur, N.S., 1978. Biostratigraphical aspects of the Subathu Formation KumaunHimalayas. Recent Researches in Geology 5, 96–112.

Mathur, A.K., Mishra, V.P., Mehra, S., 1996. Systematic study of plant fossils fromDagshai, Kasauli and Dharamsala formations of Himachal Pradesh. GeologicalSurvey of India Palaeontologia Indica (New Series) 50, 1–121.

Maynard, J.B., 1992. Chemistry of modern soils as guide to interpreting PrecambrianPaleosols. Journal of Geology 100, 279–289.

Metais, G., Antoine, P.O., Baqri, S.R.H., Crochet, J.Y., De Franceschi, D., Marivaux, L.,Welcomme, J.L., 2009. Lithofacies, depositional environments, regional biostratig-raphy and age of the Chitarwata Formation in the Bugti Hills, Balochistan, Pakistan.Journal of Asian Earth Sciences 34, 154–167.

Miller, K.G., Fairbank, R.G., Mountain, G.S., 1987. Tertiary oxygen isotope synthesis, sealevel history, and continental margin erosion. Paleoceanography 2, 1–19.

Miller, K.G., Wright, J.D., Fairbank, R.G., 1991. Unlocking the ice house: Oligocene–Miocene oxygen isotopes, eustasy, and margin erosion. Journal of GeophysicalResearch 96, 6829–6848.

Molnar, P., Tapponnier, P., 1975. Cenozoic tectonics of Asia: effects of a continentalcollision. Science 189, 419–426.

Mora, C.I., Driese, S.G., Colarusso, L.A., 1996. Middle to Late Paleozoic atmospheric CO2

levels from soil carbonate and organic matter. Science 271, 1105–1107.Mugnier, J.L., Chalaron, E., Mascle, G., Pradier, B., Herail, G., 1995. Structural and ther-

mal evolution of the Siwaliks of Western Nepal. Bulletin of the Geological Societyof Nepal 11, 171–179.

Murphy, C.P., 1986. Thin Section Preparation of Soils and Sediments. AcademicPublishers, Berkhamstead, AB (149 pp.).

Najman, Y., 2006. The sediment record of orogenesis: a review of approaches and tech-niques used in the Himalaya. Earth-Science Reviews 74, 1–72.

Najman, Y., 2007. Comment on “Biochronological continuity of the paleogene sedi-ments of the Himalayan foreland basin: Paleontological and other evidence”-Bhatia S. B. & Bhargava O.N., 2006, JAES 26,477-487. Journal of Asian Earth Sciences30, 417–421.

Najman, Y., Garzanti, E., 2000. Reconstructing early Himalayan tectonic evolution andpaleogeography from Tertiary foreland basin sedimentary rocks, northern India.Geological Society of America Bulletin 112, 435–449.

Najman, Y., Clift, P., Johnson, M.R.W., 1993. Early stages of foreland basin evolution inthe Lesser Himalaya, N. India. In: Treloar, P.J., Searle, M.P. (Eds.), Himalayan tecton-ics: Geological Society London, Special Publication, 74, pp. 541–558.

Najman, Y., Enkin, R.J., Johnson, M.R.W., Robertson, A.H.F., Baker, J., 1994.Palaeomagnetic dating of the earliest continental Himalayan foredeep sediments:implications for Himalayan evolution. Earth and Planetary Science Letters 128,713–718.

Najman, Y., Johnson, K., White, N., Olivers, G., 2004. Evolution of Himalayan forelandbasin, NW India. Basin Research 16, 1–24.

Nesbitt, H.W., Young, G.M., 1982. Early Proterozoic climates and plate motions inferredfrom major element chemistry of lutites. Nature 299, 715–717.

Nordt, L.C., Driese, S.D., 2010. New weathering index improves paleorainfall estimatesfrom Vertisols. Geology 38, 407–410.



Nordt, L.C., Orosz, M., Driese, S., Tubbs, J., 2006. Vertisol carbonate properties in relationto mean annual precipitation: implication for paleoprecipitation estimates. Journalof Geology 114, 501–510.

Pagani, M., Zachos, J., Freeman, K.H., Bohaty, S., Tipple, B., 2005. Marked change in atmo-spheric carbon dioxide concentrations during the Oligocene. Science 309, 600–603.

Pal, D.K., Balpande, S.S., Srivastava, P., 2001. Polygenetic vertisols of the Purna Valley ofCentral India. Catena 43, 231–249.

Parkash, B., Sharma, R.P., Roy, A.K., 1980. The Siwalik group (Molasse) sediments shedby collision of continental plates. Sedimentary Geology 25, 127–159.

Passey, B.J., Robinson, T.F., Ayliffe, L.K., Cerling, T.E., Spoheimer, M., Dearing, M.D.,Roeder, B.L., Ehleringer, J.R., 2005. Carbon isotope fractionation between diet,breath CO2, and bioapatite in different mammals. Journal of Archaeological Science32, 1459–1470.

Pendall, E.G., Harden, J.W., Trumbore, S.E., Chadwick, O.A., 1994. Isotopic approach to soilcarbonate dynamics and implications for paleoclimatic interpretations. QuaternaryResearch 42, 60–71.

PiPujol, M.D., Buurman, P., 1994. The distinction between groundwater gley andsurface-water gley phenomena in Tertiary paleosols of the Ebro basin, NE Spain.Palaeogeography, Palaeoclimatology, Palaeoecology 110, 103–113.

Prochnow, S., Nordt, L., Atchley, S., Hudec, M., 2006. Multi-proxy paleosol evidencefor Middle and Late Triassic climate trends in eastern Utah. Palaeogeography,Palaeoclimatology, Palaeoecology 232, 53–72.

Quade, J., Cerling, T.E., 1995. Expansion of C4 grasses in the late Miocene of northernPakistan: evidence from stable isotopes in paleosols. Palaeogeography,Palaeoclimatology, Palaeoecology 115, 91–116.

Quade, J., Cater, M.L., Ojha, J.A., Harrison, T.M., 1995. Late Miocene environmentalchange in Nepal and the nortern Indian subcontinent: stable isotopic evidence.Geological Society of America Bulletin 107, 1381–1397.

Raiverman, V., 1979. Stratigraphy and facies distribution, Subathu sediments, SimlaHills, northwestern Himalaya. Geological Survey of India Miscellaneous Publica-tions 41, 111–125.

Raiverman, V., 2002. Foreland Sedimentation in Himalayan Tectonic Regime: a Relook atthe Orogenic Process. Bishen Singh Mahendra Pal Singh, Dehra Dun, India (378 pp.).

Raiverman, V., Raman, K.S., 1971. Facies relations in Subathu sediments, Simla Hills,N.W. Himalaya, India. Geological Magazine 108, 329–341.

Raiverman, V., Suresh, N., 1997. Clay mineral distribution in the Cenozoic sequence ofthe western Himalayan Foothills. Journal of Indian Association of Sedimentologists16, 63–75.

Raiverman, V., Kunte, S.V., Mukherjea, A., 1983. Basin geometry, Cenozoic sedimenta-tion and hydrocarbon prospects in northwestern Himalayas and Indo-GangeticPlains. Petroleum Asia Journal 6, 67–92.

Ramstein, G., Fluteau, F., Besse, J., Joussaume, S., 1997. Effect of Orogeny, plate motionand land-sea distribution on Eurasian climate change over the past 30 millionyears. Nature 386, 788–795.

Ravikant, V., Wu, F.Y., Ji, W.Q., 2011. U–Pb age and Hf isotopic constraints of detrital zir-cons from the Himalayan foreland Subathu sub-basin on Tertiary palaeogeographyof the Himalaya. Earth and Planetary Science Letters 304, 356–368.

Retallack, G.J., 1991. Miocene Paleosols and Ape Habitats in Pakistan and Kenya. OxfordUniversity Press, New York (246 pp.).

Retallack, G.J., 1997a. Early forest soils and their role in Devonian global change.Science 276, 583–585.

Retallack, G.J., 1997b. Neogene expansion of North American prairie. Palaios 12,380–390.

Retallack, G.J., 1997c. Paleosols in the upper Narrabeen Group of New South Wales asevidence of early Triassic paleoenvironments without modern analogues. AustralianJournal of Earth Sciences 44, 185–201.

Retallack, G.J., 2001a. Cenozoic expansion of grasslands and climate cooling. Journal ofGeology 109, 407–426.

Retallack, G.J., 2001b. Soils of the Past. Blackwell, Oxford (600 pp.).Retallack, G.J., 2005. Pedogenic carbonate proxies for amount and seasonality of precip-

itation in paleosols. Geology 33, 333–336.Retallack, G.J., 2009. Refining a pedogenic-carbonate CO2 paleobarometer to quantify a

middle Miocene greenhouse spike. Palaeogeography, Palaeoclimatology, Palaeo-ecology 281, 57–65.

Retallack, G.J., Dilcher, D.L., 2012. Outcrop versus core and geophysical log interpretationof mid-Cretaceous paleosols from the Dakota Formation of Kansas. Palaeogeography,Palaeoclimatology, Palaeoecology 329–330, 47–63.

Retallack, G.J., Bestland, E.A., Fremd, T.J., 2000. Eocene and Oligocene paleosols of centralOregon. Geological Society of America Special Paper 344, 192pp.

Richter, F.M., Lovera, O.M., Harrison, T.M., Copeland, P., 1991. Tibetan tectonics from40Ar/39Ar analysis of a single K-feldspar sample. Earth and Planetary ScienceLetters 105, 266–278.

Rowley, D.B., Pierrehumbert, R.T., Currie, B.S., 2001. A new approach to stable isotope-basedpaleoaltimetry: implications for paleoaltimetry and paleohypsometry of the HighHimalaya since the Late Miocene. Earth and Planetary Science Letters 188, 253–268.

Sahni, A., Bhatia, S.B., Kumar, K., 1983. Faunal evidence for the withdrawal of the Tethysin the Lesser Himalaya, northwestern India. Bollettino della Societa PaleontologicaItaliana 22, 77–86.

Saini, N.K., Mukherjee, P.K., Rathi, M.S., Khanna, P.P., 2000. Evaluation of ED-XRF perfor-mance in rapid analysis of silicate rock analysis. X-Ray Spectroscopy 29, 166–172.

Saini, N.K., Mukherjee, P.K., Rathi, M.S., Khanna, P.P., Purohit, K.K., 2002. Trace elementestimation in soils: an appraisal of EDXRF technique using group analysis scheme.Journal of Trace and Microprobe Techniques 20, 539–551.

Sangode, S.J., Kumar, R., Siddaaiah, N.S., 2005. Magnetic polarity and rock magneticstudies across marine to continental transition (Subathu-Dagshai sequence)in the Himalayan Foreland. Singapore, Proceedings, Asia Oceania GeosciencesMeeting (284 pp.).

Sehgal, J.L., Stoops, G., 1972. Pedogenic calcite accumulation in arid and semi-aridregions of the Indo-Gangetic alluvial plain of erstwhile Punjab (India). Geoderma8, 59–72.

Sheldon, N.D., Tabor, N.J., 2009. Quantitative paleoenvironmental and paleoclimatic re-construction using paleosols. Earth-Science Reviews 95, 1–52.

Sheldon, N.D., Retallack, G.J., Tanaka, S., 2002. Geochemical climofunctions from NorthAmerica soils and application to paleosols across Eocene–Oligocene boundary inOregon. Journal of Geology 110, 683–696.

Sheldon, N.D., Mithchell, R.L., Collinson, M.E., Hooker, J.J., 2009. Eocene–Oligocene transi-tion record from paleosols, Isle of Wight (UK). In: Koeberl, C., Monatanri, A. (Eds.),The late Eocene Earth — hothouse, icehouse, and impacts: Geological Society ofAmerica Special Paper, 452, pp. 241–248.

Singh, B.P., 2003. Evidence of growth fault and forebulge in the late Palaeocene(57.9–54.7 Ma), western Himalayan foreland basin, India. Earth and PlanetaryScience Letters 216, 717–724.

Soil Survey Staff, 1999. Key to Soil Taxonomy, A basic System of Soil Classification formaking and Interpreting Soil Surveys, 2nd ed. Agriculture Handbook, SCS-USDA,436. US Govt. Printing Office, Washington, D.C. (869 pp.).

Soil Survey Staff, 2006. Key to Soil Taxonomy, 10th ed. USDA Natural Resources Conser-vation Service, Washington D.C.(341 pp.).

Solomon, D.K., Cerling, T.E., 1987. The annual carbon di-oxide cycle in a montane soil:observation, modelling and implications for weathering. Water Resources Research23, 2257–2265.

Srivastava, P., 2001. Paleoclimatic implications of pedogenic carbonates in Holocene soilsof the Gangetic Plains, India. Palaeogeography, Palaeoclimatology, Palaeoecology 172,207–222.

Srivastava, P., Parkash, B., 2002. Polygenetic soils of the north-central part of theGangetic Plains: a micromorphological approach. Catena 46, 243–259.

Srivastava, P., Parkash, B., Sehgal, J.L., Kumar, S., 1994. Role of neotectonics and climatein development of the Holocene geomorphology and soils of the Gangetic Plainsbetween the Ramganga and Rapti rivers. Sedimentary Geology 94, 129–151.

Srivastava, P., Bhattacharyya, T., Pal, D.K., 2002. Significance of the formation of calciumcarbonate minerals in the pedogenesis and management of cracking clay soils(Vertisols) of India. Clays and Clay Minerals 50, 111–126.

Srivastava, P., Singh, A.K., Parkash, B., Singh, A.K., Rajak, M.K., 2007. Paleoclimatic implica-tions of micromorphic features of Quaternary paleosols of NW Himalayas and poly-genetic soils of the Gangetic Plains — a comparative study. Catena 70, 169–184.

Srivastava, P., Rajak, M., Singh, L.P., 2009. Late Quaternary alluvial fans and paleosols ofthe Kangra Basin, NW Himalaya: tectonic and paleoclimatic implications. Catena76, 135–154.

Srivastava, P., Rajak, M.K., Sinha, R., Pal, D.K., Bhattacharyya, T., 2010. A high-resolutionmicromorphological record of the late Quaternary paleosols from Ganga–Yamunainterfluve: stratigraphic and paleoclimatic implications. Quaternary International227, 127–142.

Srivastava, G., Srivastava, R., Mehrotra, R.C., 2011. Ficus palaeoracemosa sp. nov. — anew fossil leaf from the Kasauli Formation of Himachal Pradesh and itspalaeoclimatic significance. Journal of Earth System Science 120, 253–262.

Srivastava, G., Spicer, R.A., Spicer, T.A.V., Yang, J., Kumar, M., Mehrotra, R., Mehrotra, N.,2012. Megaflora and paleoclimate of Late Oligocene tropical delta, Makum Coalfield,Assam: evidence for early development of South Asia Monsoon. Palaeogeography,Palaeoclimatology, Palaeoecology 342–343, 130–142.

Utescher, T., Bruch, A.A., Micheels, A., Mosbrugger, V., Popova, S., 2011. Cenozoic cli-mate gradients in Eurasia— a palaeo-perspective on future climate change?Palaeogeography, Palaeoclimatology, Palaeoecology 304, 351–358.

Warwick, P.D., Johnson, E.A., Khan, I.H., 1998. Collision-induced tectonism along thenorthwestern margin of the Indian subcontinent as recorded in the Upper Paleo-cene to Middle Eocene strata of central Pakistan (Kirthar and Sulaiman Ranges).Palaeogeography, Palaeoclimatology, Palaeoecology 142, 201–216.

Welcomme, J.L., Benammi, M., Crochet, J.Y., Mmarivaux, L., Metais, G., Antoine, P.E.,Baloch, I., 2001. Himalayan forelands: palaeontological evidence for Oligocene de-trital deposits in the Bugti Hills (Balochistan, Paksitan). Geological Magazine 138,397–405.

Wieder, M., Yaalon, D.M., 1982. Micromorphological fabrics and developmental stagesof carbonate nodular forms related to soil characteristics. Geoderma 28, 203–220.

Wilding, L.P., Tessier, D., 1988. Genesis of Vertisols: shrink-swell phenomena. In:Wilding, L.P., Punters, R. (Eds.), Vertisols: Their distribution, properties, classifica-tion and management. Texas A&M University, College Station, Texas, pp. 55–96.

Wobus, C., Heimsath, A., Whipple, K., Hodges, K., 2005. Active out of sequence thrustfaulting in Central Nepalese Himalaya. Nature 434, 1008–1011.

Zachos, J., Pagani, M., Sloan, L., Thomas, E., Billups, K., 2001. Trends, rhythms, and aber-rations in global climate 65 Ma to present. Science 292, 686–693.

Zachos, J.C., Dickens, G.R., Zeebe, R.E., 2008. An early Cenozoic perspective on green-house warming and carbon-cycle dynamics. Nature 451, 279–283.


Documents

2013 Sed Geol