Parental Affinities and Environments of Bauxite Genesis in the Salt Range, Northwestern Himalayas, Pakistan

Abstract

As the residual products of severe chemical weathering, bauxite deposits serve both as essential economic Al-Fe resources and geochemical archives that reveal information about the parent rocks’ composition, paleoenvironments and paleoclimates, and the tectonic settings responsible for their genesis. The well-developed Early Paleocene bauxite deposits of the Salt Range, Pakistan, provide an opportunity for deciphering their ore genesis and parental affinities. The deposits occur as lenticular bodies and are typically composed of three consecutive stratigraphic facies from base to top: (1) massive dark-red facies (L-1), (2) composite conglomeratic–pisolitic facies (L-2), and (3) Kaolinite-rich clayey facies (L-3). Results from optical microscopy, X-ray powder diffraction (XRPD), and scanning electron microscopy with Energy-Dispersive X-Ray Spectroscopy (SEM-EDS) reveal that facies L-1 contains kaolinite, hematite, and goethite as major minerals, with minor amounts of muscovite, quartz, anatase, and rutile. In contrast, facies L-2 primarily consists of kaolinite, boehmite, hematite, gibbsite, goethite, alunite/natroalunite, and zaherite, with anatase, rutile, and quartz as minor constituents. L-3 is dominated by kaolinite, quartz, and anatase, while hematite and goethite exist in minor concentrations. Geochemical analysis reveals elevated concentrations of ${\mathrm{Al}}_2$${\mathrm{O}}_3$, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$, ${\mathrm{SiO}}_2$, and ${\mathrm{TiO}}_2$. Trace elements, including Th, U, Ga, Y, Zr, Nb, Hf, V, and Cr, exhibit a positive trend across all sections when normalized to Upper Continental Crust (UCC) values. Field observations and analytical data suggest a polygenetic origin of these deposits. L-1 suggests in situ lateritization of some sort of precursor materials, with enrichment in stable and ultra-stable heavy minerals such as zircon, tourmaline, rutile, and monazite. This facies is mineralogically mature with bauxitic components, but lacks the typical bauxitic textures. In contrast, L-2 is texturally and mineralogically mature, characterized by various-sized pisoids and ooids within a microgranular-to-microclastic matrix. The L-3 mineralogy and texture suggest that the conditions were still favorable for bauxite formation. However, the ongoing tectonic activities and wet–dry climate cycles post-depositionally disrupted the bauxitization process. The accumulation of highly stable detrital minerals, such as zircon, rutile, tourmaline, and monazite, indicates prolonged weathering and multiple cycles of sedimentary reworking. These deposits have parental affinity with acidic-to-intermediate/-argillaceous rocks, resulting from the weathering of sediments derived from UCC sources, including cratonic sandstone and shale.

1. Introduction

Bauxites are residual deposits which provide for the primary source of aluminum and iron ore, formed through the intense chemical weathering of aluminosilicate-rich parent rocks [1,2,3,4,5,6]. Climate, vegetation cover, surface water chemistry, bedrock composition and texture, tectonics and orogeny, groundwater hydrodynamics, and local topography are among the key factors controlling bauxite formation [1,7,8,9,10,11,12]. Based on host lithology, mineralogy, and geochemistry, bauxites are classified into three major genetic categories [1,13,14]: (1) Lateritic-type bauxites are products of the in situ chemical weathering (autochthonous) of aluminosilicate rocks. These deposits are common in geologically stable areas, including passive continental margins and stable plate interiors. Most of the world’s bauxite resources belong to this category [2,3,12]. (2) Tikhvin-type bauxites are detrital (allochthonous to parallochthonous) of pre-existing lateritic bauxite erosion, reworking, and redeposition [15,16]. (3) Karst-type bauxites are developed through chemical weathering, accumulating on karstified carbonate rocks, and are rich in boehmite and diaspore. These types of bauxite are mainly found in mountainous or orogenic regions where tectonics causes uplift, subaerial exposure, and karstification, playing important roles in its formation [1,3,17,18]. These are important sources of alumina and rare earth elements (REE) and have received significant research focus [11,12,19,20,21,22,23,24,25,26,27]. According to paleogeographic models, a strong relationship exists between the tectonic and paleoclimatic conditions of the Tethyan belt bauxite deposits [14,28,29].

The Salt Range in the northwestern Himalayas, Pakistan, contains karst bauxite deposits primarily occurring as lenticular bodies within the unconformity between the Permian limestone and the Paleocene Hangu Formation [30,31]. Although the presence of bauxite deposits has been documented in the field, a comprehensive genetic model explaining their source rocks, parental affinities, and the paleoenvironmental conditions that led to their formation is still underexplored. Previous research has mainly focused on the economic importance of these deposits and the extraction of alumina. Locally, the studied bauxite deposits are currently exploited for alumina, iron oxides, and kaolinite clays [30,31,32,33,34,35].

This study aims to investigate the genetic model of bauxite deposits and their parental affinities, as such insights are essential for fundamental geological understanding and also for improving exploration strategies. A comprehensive genetic model will also provide information on the paleogeographic position of the Indian Plate and the regional tectonic framework prior to its collision. A multiproxy approach is used to identify the probable parent rocks, establish a robust genetic model, and reconstruct the paleoenvironmental and tectonic conditions that facilitate bauxitization in the region. The methodology combines detailed field observations with textural and mineralogical analysis, scanning electron microscopy, and whole-rock major and trace element.

2. Regional Geology and Tectonics of the Salt Range

The Salt Range in Pakistan, along with its extension across the Indus River, connects the Sulaiman Mountain arc to the outer regions of the northwestern Himalayas (Figure 1a). This area features an exposed stratigraphic sequence that spans from the Precambrian to the Holocene epoch [36,37]. The oldest exposed unit is the Salt Range Formation, which consists primarily of evaporites, while the youngest sediments are classified as Recent Conglomerates [37,38]. The Salt Range is subdivided into three sections: eastern, central, and western. The studied bauxite/laterite deposits are mainly located in the central Salt Range [37,39]. Structurally, the Salt Range is quite complex, and young uplift resulted from compressional deformation linked to the Himalayan orogeny. In the central Salt Range, tight anticlines are intermingled with broad synclines [37,40] (Figure 1b).

Four major unconformities shape the exposure and arrangement of the exposed strata. These regional unconformities interrupt the stratigraphic succession, indicating significant changes in tectonic activity and depositional regimes over geological time [37,38,41,42]. The first major unconformity is at the base of the Permian, where the glaciofluvial conglomerates of the Tobra Formation rest unconformably upon a Cambrian succession [37,43]. The second unconformity occurs at the base of the Cenozoic, where Paleocene strata directly overlie the Cambrian Jehlum Group and Permian strata. This unconformity is characterized by the studied bauxite/laterite deposits at places in the central and eastern parts (Figure 1b). At the third unconformity, Upper Miocene strata overlie thin Paleocene–Eocene or Triassic strata in the western section. The fourth major unconformity occurs at the base of the late Pliocene–Pleistocene on the western side of the Salt Range [37].

Mesozoic sequences are mostly absent in the central and eastern Salt Range, largely due to tectonic uplift, inadequate sediment accumulation, and subsequent erosion. These processes likely transformed the area into a paleotopographic high for most of the Mesozoic era, limiting sedimentation [37,43]. Tectonic activities during the Paleozoic–Mesozoic eras [14,42,44,45], particularly the late Cretaceous uplift of the Sargodha high [46,47] and persistent thrusting in the Salt Range, could have led to the erosion of Mesozoic rocks in the eastern and central Salt Range, leaving no remnants in the stratigraphic record [37,46,48,49,50]. However, these strata are present in the western Salt Range and Potwar Plateau, indicating that they were part of a broader depositional system. This interpretation is supported by the Paleogene bauxite and laterite beds located at the base of the Paleocene Hangu Formation in the central–eastern Salt Range, just above the prominent regional unconformity (Figure 1b) [31,37,39,42,51]. The development of these bauxites reflects the interplay of tectonic uplift, subaerial exposure, non-deposition, and the paleogeographic position of the Indian Plate, restricting their distribution to the central–eastern Salt Range.

3. Materials and Analytical Methods

This study is based on data collected through fieldwork and subsequent analytical laboratory analyses. During fieldwork, 30 samples were collected from the outcrops studied. Samples were collected on the basis of color, textural variations, and lithologies. Seven (7) samples were collected from the CMB section, four (4) from the bauxite deposit, one (1) from the underlying Permian limestone, one (1) from the overlying Paleocene Hangu sandstone, and one (1) from the Lochkart Limestone. Similarly, twenty-three (23) samples were collected from the Arara sections; twelve (12) from ATB and eleven (11) from the AKB outcrop sections, including one (01) from the underlying Permian limestone, one (1) from the overlying Paleocene Hangu sandstone, and one (01) from the Lochkart Limestone, to investigate pre- and post-depositional processes and for age confirmation (Figure 2).

X-ray powder diffraction (XRPD), optical petrography, and SEM/EDS analyses were performed on bauxite and overlying sandstone samples to identify mineralogical phases and textural complexities. These analyses were conducted at the Department of Geology, University of Vienna, Austria. For XRPD analysis, the Malvern PANalytical X‘PERT Pro diffractometer, Almelo, Netherlands (Cu K$\alpha$-radiation (40 kV, 40 mA) was used. Samples underwent qualitative XRD analysis using the Xpert Pro software, version 2.2e (2.2.5). Polished thin sections (approximately 30 $\mathsf{\mu }$$\mathrm{m}$ thick) were examined under a petrographic microscope using transmitted and reflected light. Scanning electron microscopy (SEM) was performed at the ‘Electron Beam Microanalysis’ Core Facility at the University of Vienna, Austria, using a Tescan VEGA 4GMU microscope. The elemental composition of mineral phases was determined via Energy-Dispersive Spectroscopy (EDS) using the Oxford AZteclivelite EDS. Comprehensive chemical analyses of major, minor, and trace elements, including rare earth elements (REE), were conducted on 20 bauxite samples (

Table 1. Major oxides (%) and trace and rare earth element (ppm) contents of the Salt Range bauxite. (L-1: basal massive dark-red facies, L-2: conglomeratic–pisolitic facies, L-3: Kaolinite-rich clayey facies. CMB: Chamilwala Mohar section, ATB: Arara Top Bauxite section, and AKB: Arara Khushab Bauxite section).

	L-1			L-2														L-3
Sample	CMB1	ATB1	AKB1	CMB2	CMB3	ATB2	ATB3/4	ATB5	ATB6	ATB7/1	ATB7/2	AKB2	AKB3	AKB4	AKB5	AKB6	AKB7	CMB4	ATB8	AKB8
${\mathrm{SiO}}_2$	29.09	20.75	28.70	34.93	43.31	18.48	9.21	21.43	26.35	38.14	42.71	9.06	29.97	30.38	43.25	17.34	32.52	42.31	45.27	32.31
${\mathrm{Al}}_2$${\mathrm{O}}_3$	28.09	26.18	27.42	38.69	36.50	43.43	38.92	40.67	40.38	33.45	36.57	52.81	39.25	25.54	31.13	53.75	37.53	30.56	30.30	21.78
${\mathrm{Fe}}_2$${\mathrm{O}}_3$	25.18	34.38	29.66	6.83	2.02	10.90	6.70	10.18	4.28	9.45	0.91	19.83	14.77	28.39	10.75	5.15	12.08	4.15	2.23	27.54
MgO	1.06	0.08	0.67	0.15	0.14	0.18	0.11	0.18	0.16	0.15	0.13	<0.01	<0.01	0.10	<0.01	<0.01	<0.01	0.95	0.70	0.63
CaO	0.22	0.01	0.22	0.12	0.10	0.10	0.06	0.04	0.09	0.03	0.02	0.13	0.24	0.18	0.14	0.17	0.22	0.30	0.09	0.85
${\mathrm{Na}}_2$O	0.10	0.04	0.06	0.13	0.09	0.15	0.14	0.06	1.21	0.07	0.19	0.01	0.07	0.05	0.04	1.13	0.15	0.26	0.14	0.08
${\mathrm{K}}_2$O	0.10	0.33	0.27	0.10	0.06	0.04	0.01	<0.01	0.09	0.03	0.07	0.04	0.11	0.12	0.10	0.44	0.26	2.41	2.16	1.80
${\mathrm{TiO}}_2$	1.95	2.56	1.46	3.83	2.95	3.59	3.03	3.22	3.11	2.96	3.24	3.49	2.20	1.84	2.17	3.25	2.97	2.01	3.02	2.07
${\mathrm{P}}_2$${\mathrm{O}}_5$	0.05	0.04	0.06	0.04	0.02	0.11	0.04	0.08	0.04	0.05	0.01	0.12	0.06	0.13	0.05	0.34	0.07	0.19	0.10	0.35
MnO	<0.01	0.01	0.01	<0.01	<0.01	<0.01	<0.01	<0.01	<0.01	<0.01	<0.01	0.02	0.01	<0.01	0.00	0.01	0.00	0.01	0.01	0.03
LOI	13.70	15.10	11.27	14.60	14.50	22.50	41.40	23.70	23.90	15.30	15.80	13.80	12.74	12.80	12.06	18.04	14.10	16.30	15.30	14.35
Sum	99.66	99.58	99.80	99.66	99.78	99.56	99.65	99.58	99.63	99.65	99.69	99.31	99.44	99.53	99.68	99.61	99.91	99.54	99.38	98.63
Ba	32.00	72.00	27.20	33.00	9.00	17.00	11.00	12.00	14.00	12.00	18.00	9.50	13.00	27.00	4.80	3.10	3.40	317.00	337.00	12.20
Ni	49.00	10.00	18.50	10.00	10.00	10.00	10.00	10.00	10.00	10.00	10.00	10.10	13.70	10.00	11.00	13.30	15.50	32.00	34.00	22.10
Cr	355.00	430.00	213.00	1140.00	934.00	402.00	375.00	361.00	341.00	382.00	259.00	414.70	257.70	341.00	128.80	210.70	265.10	225.00	184.00	222.50
Sc	0.20	36.00	22.10	0.30	0.10	27.00	29.00	26.00	27.00	13.00	15.00	17.30	15.70	18.00	15.50	14.70	17.50	5.90	37.00	30.10
Be	45.80	2.00	1.80	49.80	34.10	3.00	1.00	2.00	2.00	<1	<1	1.00	0.92	1.00	0.65	0.71	0.90	54.70	3.00	1.70
Co	16.70	6.80	14.40	32.70	24.10	4.90	8.40	4.70	2.90	1.60	1.00	13.50	10.30	2.40	3.50	4.10	3.90	36.10	6.90	9.50
Cs	0.20	0.40	0.78	0.30	0.10	<0.1	<0.1	<0.1	<0.1	<0.1	<0.1	<0.1	<0.1	0.30	0.20	0.27	0.16	5.90	6.00	4.20
Ga	45.80	41.40	53.00	49.80	34.10	76.90	60.50	65.70	64.80	39.60	31.90	28.10	23.40	20.60	22.30	25.60	21.20	54.70	50.40	27.50
Hf	16.70	33.80	16.20	32.70	24.10	50.30	40.90	46.10	45.70	43.50	44.20	21.50	23.10	22.70	24.20	20.50	18.20	36.10	76.80	42.50
Nb	35.40	39.30	24.50	80.20	60.20	51.90	42.50	48.30	44.90	47.50	51.20	29.50	30.20	31.50	33.20	35.60	31.50	44.20	53.00	36.30
Rb	3.60	12.20	13.50	3.10	1.80	1.30	0.40	0.30	3.30	0.90	2.50	4.40	4.10	4.30	3.80	5.10	4.60	93.00	85.40	51.50
Sn	5.00	9.00	8.20	10.00	7.00	12.00	10.00	10.00	9.00	7.00	7.00	6.70	5.60	7.00	3.70	6.50	7.50	7.00	8.00	9.30
Sr	245.00	86.50	166.50	115.50	43.00	52.00	94.10	24.60	78.10	16.20	40.40	153.20	154.50	155.80	148.80	161.50	158.10	597.70	203.00	198.50
Ta	2.60	3.10	2.20	5.10	4.00	6.10	3.30	3.70	3.70	4.30	4.60	2.60	1.90	2.40	2.70	2.40	1.80	2.90	4.70	2.30
Th	37.80	77.80	42.10	64.20	45.10	106.80	44.20	99.80	70.00	50.20	41.00	68.60	65.50	73.40	71.50	78.20	81.50	44.00	64.40	67.50
U	13.60	8.20	9.20	10.90	6.60	11.20	8.80	10.00	8.90	7.30	7.10	6.40	5.30	4.00	4.60	5.10	4.20	12.60	17.30	15.20
V	357.00	939.00	1033.00	524.00	242.00	578.00	392.00	572.00	318.00	405.00	209.00	1127.00	1295.00	1491.00	1305.00	1535.00	1275.00	305.00	216.00	1523.00
W	3.70	4.90	3.10	10.50	4.20	9.20	5.50	7.30	8.10	4.00	4.10	2.63	2.45	2.50	2.30	3.50	2.70	4.30	7.50	6.20
Zr	637.10	1244.20	565.00	1234.60	953.70	1867.20	1553.70	1775.80	1721.90	1564.60	1650.40	1873.00	865.00	893.20	1046.00	2150.00	1619.00	1423.60	2946.20	2401.00
Y	73.10	43.30	33.60	56.60	40.00	52.60	41.20	53.40	44.50	42.10	45.60	31.50	28.20	29.60	32.10	35.10	30.50	67.10	113.50	47.50
La	49.70	19.10	20.50	17.40	10.00	6.50	5.60	4.10	8.60	3.60	6.90	13.20	11.80	10.50	9.70	12.10	11.40	114.00	126.50	72.00
Ce	607.50	41.40	45.10	51.00	29.40	25.90	19.20	14.30	32.50	9.40	15.70	30.10	26.10	22.90	21.40	28.10	23.10	260.80	261.20	151.80
Pr	12.27	3.28	4.70	3.90	2.08	2.44	1.45	1.26	2.40	0.96	1.47	3.10	2.70	1.97	1.80	2.70	2.20	26.45	27.69	16.50
Nd	51.50	10.80	11.80	14.10	7.50	9.60	5.40	4.80	8.20	4.00	5.50	9.60	7.30	6.30	5.80	7.80	6.40	90.00	96.40	67.34
Sm	13.46	2.69	2.40	3.83	2.04	2.24	1.72	1.68	1.92	1.50	1.80	2.10	1.70	1.54	1.40	1.75	1.20	14.65	18.52	11.10
Eu	2.86	0.68	0.61	0.81	0.48	0.56	0.43	0.47	0.46	0.43	0.46	0.53	0.37	0.36	0.28	0.41	0.32	2.43	3.21	2.49
Gd	14.87	4.24	3.20	4.94	3.06	4.02	3.14	3.78	3.31	3.09	3.28	3.10	2.20	2.33	1.90	2.20	1.64	10.83	16.84	11.10
Tb	2.11	1.00	0.75	1.04	0.73	1.00	0.79	0.99	0.79	0.75	0.80	0.71	0.63	0.55	0.43	0.52	0.40	1.76	2.85	1.80
Dy	11.84	7.33	4.80	7.79	5.53	8.06	6.00	7.93	6.31	6.10	6.46	5.30	4.30	4.20	3.30	3.90	3.30	11.23	18.57	14.20
Ho	2.41	1.63	1.45	1.78	1.26	1.95	1.44	1.92	1.55	1.45	1.58	1.30	1.14	1.05	1.01	1.08	0.90	2.40	4.08	3.20
Er	7.25	5.48	4.60	5.84	4.07	6.78	5.03	6.68	5.44	5.06	5.50	4.80	3.71	3.57	2.80	3.50	2.80	7.57	13.50	8.20
Tm	0.98	0.84	0.70	0.90	0.63	1.06	0.76	1.01	0.82	0.80	0.85	0.74	0.68	0.55	0.48	0.64	0.48	1.21	2.04	1.40
Yb	6.51	5.90	4.10	6.31	4.38	7.66	5.74	7.58	6.15	5.75	6.36	5.60	4.60	3.92	3.20	4.60	3.70	8.35	14.75	8.40
Lu	0.92	0.88	0.58	0.95	0.67	1.17	0.83	1.13	0.92	0.89	0.98	0.80	0.75	0.62	0.51	0.72	0.58	1.28	2.26	1.40
$\Sigma$LREE	737.29	77.95	85.11	91.04	51.50	47.24	33.80	26.61	54.08	19.89	31.83	58.63	49.97	43.57	40.38	52.86	44.62	508.33	533.52	321.23
$\Sigma$HREE	46.89	27.30	20.58	29.55	20.33	31.70	23.73	31.02	25.29	23.89	25.81	22.35	18.01	16.79	13.63	17.16	13.80	44.63	74.89	49.70
$\Sigma$REE	784.18	105.25	105.69	120.59	71.83	78.94	57.53	57.63	79.37	43.78	57.64	80.98	67.98	60.36	54.01	70.02	58.42	552.96	608.41	370.93
La/Y	0.68	0.44	0.61	0.31	0.25	0.12	0.14	0.08	0.19	0.09	0.15	0.42	0.42	0.35	0.30	0.34	0.37	1.70	1.11	1.52
(La/Yb)N	5.29	2.24	3.46	1.91	1.58	0.59	0.68	0.37	0.97	0.43	0.75	1.63	1.78	1.85	2.10	1.82	2.13	9.45	5.94	5.93
Sm/Nd	0.26	0.25	0.20	0.27	0.27	0.23	0.32	0.35	0.23	0.38	0.33	0.22	0.23	0.24	0.24	0.22	0.19	0.16	0.19	0.16
Ce/Ce*	5.90	1.25	1.10	1.48	1.55	1.56	1.62	1.51	1.71	1.21	1.18	1.13	1.11	1.21	1.23	1.18	1.11	1.14	1.06	1.06
Eu/Eu*	0.61	0.61	0.67	0.57	0.58	0.57	0.56	0.57	0.55	0.61	0.57	0.63	0.58	0.58	0.52	0.63	0.69	0.59	0.55	0.68

) using Inductively Coupled Plasma Optical Emission Spectroscopy and Mass Spectrometry (ICP-OES/MS) at Bureau Veritas Commodities Canada Ltd., Vancouver, BC, Canada. Each 0.5 g sample was digested with aqua regia, followed by ${\mathrm{LiBO}}_2$ fusion. Major oxides and minor elements were quantified using ICP-OES, while REE and trace elements were determined by ICP-MS. Loss on ignition (LOI) was calculated by measuring the weight loss after heating the samples to 1000 °C.

4. Results

4.1. Outcrop Description

Three outcrops were selected for sampling, one from Chamilwala Mohar (CMB) and two from Arara: the Arara Top Bauxite (ATB) and Arara Khushab Bauxite (AKB) (Figure 1b). These deposits are complex and heterogeneous in color, shape, size, and composition (Figure 2). All three sections (CMB, ATB, and AKB) exhibit three distinct facies: a basal dark-red massive lithofacies (L-1) with a thickness varying from 0.5 to 1 m, overlain by a conglomeratic–pisolitic lithofacies (L-2) (0.5 to 2.5 m thick) and capped by a laminated clayey lithofacies (L-3) (0.5 m thick) (Figure 2 and Figure 3). The L-1 and L-3 lithofacies are largely similar across all three outcrops, while the L-2 lithofacies shows variations. The L-1 lacks internal stratification and is rich in iron oxides (Figure 2 and Figure 3a,b). The L-2 facies typically shows nodular, oolitic/pisolitic textures (Figure 2 and Figure 3c,e,f). In the CMB section, L-2 is 0.5 m thick and creamy/maroon to pale yellow in color and contains various-sized ooids, pisoids, and micropebbles (0.5 to 2 cm in diameter) (Figure 2 and Figure 3a–c). In the ATB section, it comprises multiple layers of pale-yellow-to-creamy colors with pisoids/ooids and bauxitic pebbles (0.5 to 2.5 cm in diameter), often featuring veins filled with alunite and kaolinite (Figure 2 and Figure 3d–f). In the AKB section, L-2 appears as dark-brown/yellow-to-creamy/maroon, oolitic/pisolitic layers with microclasts (0.5 to 2 cm in diameter) (Figure 2 and Figure 3g–i). Facies L-3 mainly comprises gray-colored laminated clayey lithofacies (Figure 3i).

4.2. Mineralogy and Texture

XRPD results show that the L-1 facies of the Chamilwala Mohar (CMB) and Arara (ATB and AKB) sections primarily consist of kaolinite, hematite, and goethite with minor amounts of anatase, boehmite, mica, quartz, and gibbsite (Figure 2 and Figure 4). The overlying L-2 facies predominantly contain kaolinite and bauxite minerals such as boehmite, gibbsite, and alunite/natroalunite and minor quantities of anatase, rutile, mica, and quartz. Moreover, an unusual mineral, zaherite, is reported from the pale-yellow L-2 of the Arara (ATB) section (Figure 2 and Figure 4). The L-3 facies consists of kaolinite and quartz as major minerals and mica and goethite as minor constituents (Figure 2 and Figure 4).

Polished thin sections and SEM/EDS results reveal that the textural components identified in these bauxite ores include concentric features such as micro-ooids, ooids, pisoids, macropisoids, nodules, concretions, microclasts (fragments), detrital grains, fissure fillings, and secondary films and coatings. The texture of the L-1 facies (Figure 5a,b) differs from that of the overlying L-2 facies (Figure 5c–e).

L-1 is rich in iron nodules with gibbsite/kaolinite coatings, though kaolinite nodules with iron coatings are also occasionally observed (Figure 5a,b and Figure 6b,b1,b2). This facies is enriched with detrital minerals, including monazite, zircon, rutile, tourmaline, mica, and quartz. L-2 exhibits not only color variations but also a high degree of textural complexity. This bauxite ore predominantly displays a matrix-supported texture, with grain-supported textures occurring less frequently. The structure is mainly pisolitic–conglomeratic and occasionally oolithic, featuring intralclasts of earlier bauxitic/lateritic crust and extraclasts of quartz, mica, and ultra-stable detrital minerals such as zircon, rutile, tourmaline, and monazite (Figure 5c–e and Figure 6e,f). Pisoids are more abundant than ooids, showing significant compositional variability, with hematitic cores and boehmite/kaolinitic rims, and, in some cases, the reverse (Figure 5c–e and Figure 6a–d). Some ooids and pisoids also exhibit partial or complete rim replacement by gibbsite/boehmite (Figure 5c,d and Figure 6b–d). In many samples, micropebbles of bauxites are observed within a whitish, clay-rich kaolinitic and boehmitic matrix and microoids and ooids embedded in the pisoids cores (Figure 5e). The matrix displays considerable color variation, ranging from dark brown to whitish-yellow and maroon (Figure 5c–e). This facies contains a wide range of fine-grained detrital minerals, including monazite, zircon, mica, quartz, and rutile (Figure 5f–i and Figure 6e,f). The bauxite bodies are capped by L-3 facies, which is enriched in fine-grained siliciclastic materials and exhibits a microclastic texture, containing detrital minerals such as zircon, tourmaline, mica, quartz and rutile (Figure 5f–h and Figure 6e,f).

4.3. Major Elements Geochemistry

The geochemical data indicate that ${\mathrm{Al}}_2$${\mathrm{O}}_3$, ${\mathrm{SiO}}_2$, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$, and ${\mathrm{TiO}}_2$ are the primary components of the bauxites in the outcrop studied (CMB, ATB, and AKB). Among these, ${\mathrm{Al}}_2$${\mathrm{O}}_3$ is the most abundant constituent. In L-1, its average concentration is 27.2%, with the lowest value recorded in the AKB section and the highest in the CMB section. L-2 exhibits fluctuating ${\mathrm{Al}}_2$${\mathrm{O}}_3$ concentrations, averaging 37.6% in CMB, 38.9% in ATB, and 40% in AKB. In L-3, the average ${\mathrm{Al}}_2$${\mathrm{O}}_3$ concentration is again lower, with a mean of 27.6% (Table 1).

Silica (${\mathrm{SiO}}_2$) is the second most abundant component, reaching its highest concentration in L-3 with an average of 40%. In L-1, the average ${\mathrm{SiO}}_2$ content is 26.2%. L-2 in CMB records an average of 39.1%, while ATB and AKB report 26.1% and 27.1%, respectively. In L-3, the ${\mathrm{SiO}}_2$ content ranges from 32.3% in AKB to 45.3% in ATB (Table 1).

${\mathrm{Fe}}_2$${\mathrm{O}}_3$ ranks third in abundance. Its highest concentration occurs in L-1, with an average of 29.7%. In L-2, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$ averages 4.4% in CMB, 7.1% in ATB, and 15.2% in AKB. L-3 shows an average of 11.3% ${\mathrm{Fe}}_2$${\mathrm{O}}_3$ (Table 1). ${\mathrm{TiO}}_2$ shows an average concentration of 2% in L-1 and 3.4% in CMB, 3.2% in ATB, and 2.6% in the AKB L-2, whereas L-3 records an average of 2.4% ${\mathrm{TiO}}_2$ (Table 1). Loss on Ignition (LOI) is another important parameter, reflecting volatile components (mainly ${\mathrm{H}}_2$O). L-1 shows an average LOI of 13.4%, while L-2 records values of 14.5% in CMB, 23.8% in ATB, and 13.9% in AKB. Notably, layers enriched with zaherite and alunite/natroalunite display comparatively higher LOI values (Table 1).

4.4. Trace Elements Geochemistry

The UCC-normalized trace element distribution pattern [52] reveals enrichment peaks for Ga, Th, U, Y, Zr, Nb, Hf, V, and Cr in all three sections (Table 1 and Figure 7a). In L-1, Zr averages 815 ppm (565–1244 ppm), Cr 332 ppm (213–430 ppm), Sr 166 ppm (86–166 ppm), and V 776 ppm (357–1033 ppm). L-2 shows elevated values, with Zr averaging 1483 ppm (865–1867 ppm), Cr 415 ppm (128–1140 ppm), Sr 99 ppm (16–161 ppm), and V 804 ppm (209–1535 ppm). In L-3, Zr further increases to an average of 2256 ppm (1423–2946 ppm), while Cr averages 210 ppm (184–222 ppm), Sr 333 ppm (203–597 ppm), and V 681 ppm (216–1523 ppm) (Table 1). Among the individual sections, ATB records markedly higher Cr (184–430 ppm), V (209–939 ppm), and Zr (1244–2946 ppm) concentrations, while other elements remain within lower ranges. In the AKB section, Cr varies from 213 to 414 ppm, V from 1033 to 1535 ppm, Zr from 565 to 2401 ppm, and Sr from 149 to 198 ppm, with other elements again restricted to a few or tens of ppm (Table 1 and Figure 3).

The major oxides and bauxitophile elements, such as Ga, Nb, Nd, Gd, Cr, V, and Co, show uneven correlation patterns. They either correlate weakly positively or show negative correlations. ${\mathrm{Al}}_2$${\mathrm{O}}_3$ has a strong positive correlation only with ${\mathrm{TiO}}_2$. ${\mathrm{TiO}}_2$ shows a strong positive correlation with ${\mathrm{Al}}_2$${\mathrm{O}}_3$, Nb, and Zr and a strong negative correlation with ${\mathrm{Fe}}_2$${\mathrm{O}}_3$. Moreover, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$ shows a strong negative correlation with ${\mathrm{Al}}_2$${\mathrm{O}}_3$, ${\mathrm{TiO}}_2$, and Nb, but it has a strong positive correlation with V (

Table 2. Correlation table: red = negative correlation, green = positive correlation.

	${\mathrm{SiO}}_2$	${\mathrm{Al}}_2$${\mathrm{O}}_3$	${\mathrm{Fe}}_2$${\mathrm{O}}_3$	CaO	${\mathrm{Na}}_2$O	${\mathrm{K}}_2$O	${\mathrm{TiO}}_2$	LOI	Ga	Nb	Nd	Gd	Cr	Zr	V	Ni	Sc	Co
${\mathrm{SiO}}_2$	1.00
${\mathrm{Al}}_2$${\mathrm{O}}_3$	−0.51	1.00
${\mathrm{Fe}}_2$${\mathrm{O}}_3$	−0.31	−0.49	1.00
CaO	0.11	−0.39	0.36	1.00
${\mathrm{Na}}_2$O	−0.18	0.44	−0.38	−0.07	1.00
${\mathrm{K}}_2$O	0.40	−0.38	−0.09	0.53	0.01	1.00
${\mathrm{TiO}}_2$	−0.26	0.72	−0.59	−0.46	0.25	−0.27	1.00
LOI	−0.53	0.30	−0.34	−0.25	0.24	−0.15	0.37	1.00
Ga	−0.24	0.08	−0.22	−0.29	0.14	0.04	0.27	0.58	1.00
Nb	0.31	0.11	−0.57	−0.29	0.01	0.03	0.63	0.18	0.41	1.00
Nd	0.39	−0.43	−0.04	0.48	−0.07	0.93	−0.29	−0.16	0.14	0.07	1.00
Gd	0.31	−0.44	0.04	0.36	−0.12	0.73	−0.22	−0.12	0.22	0.14	0.92	1.00
Cr	0.04	0.15	−0.17	−0.23	−0.13	−0.31	0.46	0.01	0.15	0.75	−0.24	−0.14	1.00
Zr	−0.07	0.26	−0.37	0.16	0.28	0.50	0.50	0.25	0.13	0.20	0.40	0.33	−0.24	1.00
V	−0.26	0.01	0.54	0.48	−0.01	−0.07	−0.34	−0.35	−0.65	−0.64	−0.23	−0.34	−0.33	−0.06	1.00
Ni	0.28	−0.40	0.12	0.34	−0.09	0.56	−0.43	−0.21	0.10	−0.12	0.80	0.89	−0.26	0.02	−0.24	1.00
Sc	−0.34	−0.12	0.23	0.03	0.05	0.20	0.03	0.35	0.26	−0.25	0.08	0.06	−0.48	0.54	0.15	−0.18	1.00
Co	0.22	−0.08	−0.10	0.18	−0.17	0.34	−0.08	−0.16	0.16	0.40	0.40	0.32	0.55	−0.28	−0.29	0.34	−0.59	1.00

).

4.5. Rare Earth Elements Geochemistry

The rare earth element (REE) content is normalized to chondrite values [53], and the resulting distribution patterns are presented in Figure 7b. Among all facies studied (CMB, ATB, and AKB), the L-3 facies exhibits the highest total REE concentration ($\Sigma$REE: La–Lu) at 510 ppm, followed by the L-1 facies with 331 ppm and the L-2 facies with 68 ppm. Light rare earth elements (LREEs: La–Sm) are predominant across all three facies, with concentrations of 300 ppm in L-1, 46 ppm in L-2, and 454 ppm in L-3. In contrast, heavy rare earth elements (HREEs: Eu–Lu) show significantly lower concentrations: 31 ppm in L-1, 22 ppm in L-2, and 56 ppm in L-3 (Table 1).

The REE fractionation in the bauxite profile is characterized by the La/Y and (La/Yb)N ratios. The La/Y ratios range from 0.44 to 0.68 in L-1, from 0.08 to 0.37 in L-2, and from 1.11 to 1.70 in L-3. Correspondingly, the (La/Yb)N ratios vary from 2.24 to 5.29 in L-1, from 0.37 to 2.13 in L-2, and from 5.93 to 9.45 in L-3, indicating greater LREE enrichment in the upper and basal facies. Additional geochemical proxies include Cerium (Ce/Ce*) and Europium (Eu/Eu*) anomalies. Ce/Ce* values range from 1.10 to 5.9 in L-1, from 1.11 to 1.71 in L-2, and from 1.06 to 1.14 in L-3, while Eu/Eu* values range narrowly between 0.55 and 0.69 across all facies (Table 1 and Figure 9a–c).

5. Discussion

5.1. Bauxite Types and Stratigraphy

Based on their stratigraphic relationship with the underlying limestones, the bauxite deposits are classified as karst-type bauxite deposits [1]. Lenticular beds of bauxite/bauxitic clay unconformably overlie the Permian Wargal Limestone and are overlain by medium-to-thick bedded yellowish sandstones of the Paleocene Hangu Formation, confirming the Paleocene age of the studied bauxites (Figure 2). The Lockhart Limestone and the Patala Formation represent subsequent marine Paleocene strata [31,54]. The bauxite outcrops are repeated due to multiple parallel sets of anticlines and synclines in the study area [37,40] (Figure 1b).

The fine-grained texture and lack of internal stratification in L-1 are indications of in situ lateritization [55]. Meanwhile, the deep-red hue in the basal part indicates ferralitic weathering under oxidation conditions [14,55]. The ferruginous nodules and ooids/pisoids result from tropical weathering conditions [1,18,55]. In L-2, pisolitic–conglomeratic textures confirm this phenomenon. The changes in the color and texture of the overlying L-3 facies indicate shifts in depositional environments.

5.2. Mineralogy and Texture

In L-1, hematite emerges as the dominant Fe-bearing phase, typically forming in oxidizing environments with a pH between 7 and 8. Ferromagnesian minerals dissolve under oxidizing conditions controlled by water activity, producing iron-rich minerals such as hematite and goethite [56,57].

The occurrence of boehmite in L-2 signifies the transformation of kaolinite into aluminum hydroxide, a process facilitated during weathering and diagenesis. Efficient groundwater circulation (i.e., good drainage) under humid tropical conditions is necessary for the removal of excess silica [58]. Gibbsite, apparently forming during the late stages of bauxitization, often acts as a cementing agent, filling desiccation cracks in ooids and pisoids. Its formation results from the ferralitic weathering of clays and K-feldspar minerals, and its fine grain size suggests precipitation from Al-rich solutions during weathering [2]. The abundance of kaolinite and boehmite reflects moderate-to-intense weathering conditions [1,59]. Additionally, the low concentration of quartz in the bauxite horizons supports the occurrence of intense chemical weathering during the dissolution phase [1].

Typically, bauxite lacks sulfate minerals such as alunite/natroalunite, which require specific conditions for formation, including a sulfur source (e.g., pyrite or volcanic gases), acidic environments during weathering, and available “${\mathrm{K}}^{+}$” and “${\mathrm{Na}}^{+}$” [60,61,62]. The Palaeocene Hangu Formation in Hazara, Trans-Indus Ranges, and the western Salt Range contain pyrite-rich shales and coal seams overlying the studied bauxites [63]. In this context, the formation of alunite–natroalunite and zaherite is interpreted as a post-depositional (late-diagenetic-to-telogenetic) phenomenon. The proposed mechanism involves pyrite oxidation in the overlying coal beds and black shales, generating sulfuric acid and acidifying percolating water. The alteration of Na-bearing minerals in these sediments likely increased sodium and potassium concentrations in the infiltrating waters. As this originally acidic, and, with time, gradually more and more alkali-enriched, meteoric water descended through channels, cracks, and pores into the underlying bauxite deposits, it reacted with the kaolinite, precipitating alunite/natroalunite. Over time, the acidic water became dispersed laterally, promoting further reactions with kaolinite [64]. Zaherite is a hydrated aluminum sulfate mineral which is extremely rare in bauxite deposits because it requires arid-to-semi-arid climates and an acidic environment rich in sulfate [65]. In contrast, bauxite formation typically occurs under humid tropical-to-subtropical conditions with high rainfall and intense ferralitic weathering [1]. If present, zaherite is restricted to specific microenvironments such as protected cavities or late fracture-filling materials [65]. In total accordance with this, in the sections studied, zaherite was observed within fractures and veins, and as a circumgranular fracture fill around the rims of the pisoids (Figure 6d).

A comparatively higher hematite concentration in the basal portions of the bauxite profiles reflects multiple stages of groundwater-related mobilization and redistribution [2,66]. Furthermore, the presence of anatase indicates residual enrichment resulting from chemical weathering and leaching [1]. L-1 also contains significant concentrations of detrital minerals such as zircon, tourmaline, rutile, and monazite. These are largely inherited from the original geochemical composition of the precursor material of the bauxite that was later modified by intense chemical weathering [67]. Chemically resistant minerals like monazite persist in these basal portions, reflecting their poor solubility in low-pH weathering conditions and at ambient temperatures [68].

The occurrence of Fe-rich nodules and kaolinite/boehmite coatings in L-1 suggests in situ weathering [1]. Fluctuations of the groundwater table appear to have played a significant role in the formation and evolution of these iron nodules [55]. The groundwater and rainwater percolating through the rock leached silica, alkali, and alkaline earth elements, leading to enrichment in Fe-bearing nodules, hematite, goethite, and kaolinite in L-1. Secondary coatings on hematite were probably formed during diagenetic or epigenetic processes (Figure 5a,b and Figure 6a,b) [2].

L-2 exhibits angular-to-subangular, often fragmented, subhedral crystal forms, indicating a degree of transport and, therefore, a detrital origin [69]. According to Bárdossy [1], the presence of round grain textures may support potentially an at least parallochthonous origin for these ores. Secondary coatings likely formed during diagenetic processes (Figure 5c–e) [2]. L-2, with oolitic/pisolitic-to-microgranular textures, appears to result from the weathering and reworking of siliciclastic and weakly bauxitized materials [1,70]. These features suggest development under intense chemical weathering, tropical climate conditions, more or less intense local parallochthonous transport by probably torrential rains, and fluctuating groundwater levels [2,18,55,70]. Moreover, the presence of fractured pisoids and hematite clasts, along with the elongated textural components, the occurrence of stretched ooids with hematite cores, and the presence of elliptical shapes, micro-ooids, and ooids within the cores of pisoids, further supports the multi-cyclic allochthonous-to-parallochthonous interpretation of these facies influenced by tectonic activities (Figure 5c–e) [71].

The presence of kaolinite, goethite, and anatase in the L-3 facies suggests that, towards the end of the apparent stratigraphic gap, climatic conditions were still favorable for ferralitic weathering. However, ongoing tectonic activities prevented the facies from maturing both texturally and mineralogically [1,48,72]. The accumulation of highly stable detrital minerals, such as zircon, rutile, tourmaline, and monazite, indicates prolonged weathering and multiple cycles of sedimentary reworking (Figure 5f–i) [73].

5.3. Major Elements Geochemistry

The UCC-normalized geochemical patterns [52] (Table 1 and Figure 7a) highlight elevated levels of ${\mathrm{Al}}_2$${\mathrm{O}}_3$, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$, and ${\mathrm{TiO}}_2$ and moderate concentrations of residual stable minerals, which generally reflect intense chemical weathering and leaching processes [15,74]. Most of the samples analyzed fall short of economy-grade bauxite criteria, except for two layers within L-2, AKB-2, and AKB-6 of the AKB section (Table 1 and Figure 2).

Correlation analysis reveals uneven, weak negative relationships between ${\mathrm{Al}}_2$${\mathrm{O}}_3$, ${\mathrm{Fe}}_2$${\mathrm{O}}_3$, ${\mathrm{TiO}}_2$, LOI, Ga, Zr, and V and ${\mathrm{SiO}}_2$ (Table 2), indicating differential silica depletion during lateritization and progressive weathering [8]. Conversely, ${\mathrm{Al}}_2$${\mathrm{O}}_3$ shows a strong positive correlation with ${\mathrm{TiO}}_2$ and a weak positive association with ${\mathrm{Na}}_2$O, LOI, Nb, Cr, and Zr (Table 2), suggesting residual enrichment of these elements during weathering. Furthermore, the weak negative correlation of ${\mathrm{K}}_2$O with ${\mathrm{TiO}}_2$ points to selective potassium leaching during bauxitization (Table 2) [13,20]. No significant correlation is observed between ${\mathrm{Al}}_2$${\mathrm{O}}_3$ and Ga, while a weak positive correlation is observed with Zr (Table 2).

The ternary diagram of major oxides (${\mathrm{SiO}}_2$–${\mathrm{Al}}_2$${\mathrm{O}}_3$–${\mathrm{Fe}}_2$${\mathrm{O}}_3$) [57] suggests that the studied bauxites experienced only weak-to-moderate chemical weathering in all three sections (CMB, ATB, and AKB) (Figure 8a). Similarly, the ${\mathrm{Fe}}_2$${\mathrm{O}}_3$–${\mathrm{Al}}_2$${\mathrm{O}}_3$–${\mathrm{SiO}}_2$ ternary diagram [1,47] indicates that most of the samples are plotted within the bauxitic clay and clayey bauxite fields, with only a few falling within the bauxite ore field (Figure 8b).

5.4. Trace Elements Geochemistry

The UCC-normalized pattern [52] (Figure 7a) further indicates that lateritization processes preferentially enriched High-Field-Strength Elements (HFSE) such as Th, U, Zr, Nb, Hf, Ti, and Ta, while significantly depleting Large-Ion Lithophile Elements (LILE), including Ba, Rb, K, Cs, Pb, Na, and Sr. Transitional Trace Elements (TTE) like Cr exhibit enrichment trends, whereas Ni and Co tend to decrease. The observed positive correlation between ${\mathrm{Al}}_2$${\mathrm{O}}_3$ and HFSE reinforces the role of aluminum in promoting the loss of TTE and more labile LILE during the lateritization process [13]. L-1 displays relatively higher ${\mathrm{Fe}}_2$${\mathrm{O}}_3$ concentrations across all three studied sections (Figure 7a). Under alkaline conditions (pH > 8), silica becomes less stable than alumina and iron oxides, leading to its leaching from phyllosilicate minerals [15,75]. Alternating wet and dry climatic cycles, fluctuations in the water table, and changing pH levels contribute to intense chemical weathering, desilicification, and soil profile development [2]. Evaporation during dry periods concentrates solutes in pore waters, increasing alkalinity, while, under wet or ferrallitic weathering conditions, acidic pore waters promote the disintegration of silicate minerals [76]. Percolating waters progressively buffer the pH from acidic in the upper horizons to alkaline at depth, generating characteristic zonation patterns in ferrallitic soils, with kaolinite and gibbsite concentrated in the upper layers and hematite accumulating in lower zones under favorable redox conditions [2,4,77,78]. The presence of boehmite further indicates advanced stages of ferrallitic weathering [2].

5.5. Rare Earth Elements Geochemistry

The bauxitization process plays a key role in REE and trace element fractionation. Factors such as the mineralogy of the parent rock, deposit geometry, degree and duration of weathering, drainage, topography, groundwater regime, microbial activity, and, most critically, climatic conditions, all influence REE concentration in bauxite [1,79]. The distribution of REEs is further affected by the presence of authigenic REE-bearing minerals and their interaction with clay phases. Resistant minerals like monazite and the adsorption of REEs onto kaolinite contribute to REE enrichment [80]. This phenomenon is particularly recorded in the L-1 and L-3 facies of the Salt Range bauxite profile, which are rich in both monazite and kaolinite (Figure 7b). In contrast, L-2 exhibits lower total REE (TREE) concentrations (Figure 9 and Table 1), suggesting more acidic weathering conditions that facilitate REE mobility and leaching during reworking [13,81]. The higher concentrations of TREE in L-1 and L-3 suggest that weathering and early diagenesis occurred under more neutral-to-alkaline pH conditions. This environment facilitates the retention of REEs through adsorption onto mineral surfaces. Furthermore, the dominance of LREEs over HREEs in all facies aligns with the lower mobility of LREEs and the presence of kaolinite and monazite [24,82].

The La/Y ratio serves as a proxy for paleoenvironmental conditions. Values below 1 in L-1 and L-2 indicate acidic conditions, while values exceeding 1 in L-3 reflect a transition to more alkaline environments [70,83]. The observed (La/Yb)N > 1.5 in L-1 and L-3 suggests typical LREE enrichment in lateritic profiles (Figure 9 and Table 1) [2,84]. These facies exhibit moderate REE fractionation associated with advanced ferralitic weathering, likely driven by high kaolinite and gibbsite contents. In contrast, the more variable and generally lower (La/Yb)N values in L-2 (Figure 9 and Table 1) suggest weaker fractionation, possibly due to less developed weathering profiles, suboptimal drainage, and reducing redox potential [85].

The Ce anomaly (Ce/Ce*) provides insight into redox conditions during weathering. Due to its low ionization potential, Ce behaves differently from other REEs [70].

In most samples, Ce/Ce* remains below 2, except for in L-1 in the CMB section, which recorded a value of 5.9, indicative of localized monazite enrichment (Figure 9 and Table 1). Cerium tends to be fixed under acidic conditions and removed in alkaline settings [15,19]. The unity values above reflect the fractionation of Ce due to the oxidation of ${\mathrm{Ce}}^{3+}$ to ${\mathrm{Ce}}^{4+}$ during intense oxidative weathering in acidic environments [86]. In addition, the strong negative Eu anomaly in all three sections indicates highly oxidizing conditions and intense weathering [87,88,89] (Figure 9 and Table 1).

5.6. Parental Affinity

Identifying the parent rocks of bauxite deposits is inherently difficult due to intense surficial weathering, erosion, and sediment transport and reworking processes [90,91,92]. Lateritic bauxites are often directly linked to their underlying source rocks through texture and composition, while this relationship is rarely evident in karst bauxites developed above carbonates. Despite considerable geological and geochemical investigations, the precise identification of source rocks and the understanding of ore-forming mechanisms in karst-type bauxites remain strongly debated [1,2]. Recently, single-grain geochronology of selected detrital mineral grains (e.g., zircons) combined with trace element geochemistry has proven to be an efficient tool in identifying the most probable source terrains of the material of some karst bauxites in the western part of the Tethyan realm [93,94,95].

In the present study, multiple geochemical proxies, including Eu/Eu* vs. Sm/Nd, Eu/Eu* vs. ${\mathrm{TiO}}_2$/${\mathrm{Al}}_2$${\mathrm{O}}_3$, and binary diagrams such as LogNi vs. LogCr [96], along with a ternary plot of Zr-Cr-Ga [97], were used to determine the parental affinities of the bauxite ore.

The Eu/Eu* vs. Sm/Nd and Eu/Eu* vs. ${\mathrm{TiO}}_2$/${\mathrm{Al}}_2$${\mathrm{O}}_3$ diagrams indicate that these sediments originated from the UCC and cratonic sandstones (Figure 10a,b). Similarly, the LogNi vs. LogCr binary diagram (Figure 11c) shows that over 80% of the deposits are classified as being formed originally as lateritic bauxite [96], supporting their derivation from sandstone and shale rather than from the lateritization of granitic, basaltic, or ultramafic rocks. This interpretation is further reinforced by the Zr-Cr-Ga ternary plot, which reveals that 100% of ATB and AKB samples, and 50% of those from CMB, were derived from acidic and intermediate/argillaceous source rocks (Figure 10d).

5.7. Ore Genesis

During the Mesozoic era, the Salt Range was located at the northwestern edge of the Indian Plate, forming part of the southwestern passive margin of the Neo-Tethys. The Late Triassic reactivation of the passive margin was marked by tectonic uplift and exhumation, which accelerated erosion across the Salt Range and adjacent areas [14,37,42,44,45,102,103]. The Tethys ocean regressed from the southeastern parts during the Late Cretaceous to Early Paleocene, likely due to the rise of the Precambrian Sargodha high to the south, exposing the southeastern region [37,46,104]. Tectonic activities such as the reactivation of the passive margin of the Indian Plate during the Late Triassic period [14,42,44,45], and particularly the rise of the Sargodha high during the Late Cretaceous [49,50,104], contributed to the subaerial exposure and erosion of Mesozoic rocks in the eastern part of the Salt Range. As a result, these rock units are absent from the stratigraphic record [37,46,51,102,105].

Figure 11. Schematic model showing formation of Salt Range bauxites: (a) Paleogeographic position of Indian Plate (Late Cretaceous–Early Paleocene) (modified from [106,107]). (b) Regional plate tectonic scenario during the Late Cretaceous–Early Paleocene (modified from [108,109,110,111]). (c) Bauxite facies formation and deposition.

Figure 11. Schematic model showing formation of Salt Range bauxites: (a) Paleogeographic position of Indian Plate (Late Cretaceous–Early Paleocene) (modified from [106,107]). (b) Regional plate tectonic scenario during the Late Cretaceous–Early Paleocene (modified from [108,109,110,111]). (c) Bauxite facies formation and deposition.

During the Late Cretaceous to Early Paleocene, persistent uplift, subaerial exposure, erosion, and weathering formed the studied bauxites. This was followed by widespread subsidence, leading to marine transgression and deposition of the Hangu Formation, the first Paleocene unit of the Salt Range.

These deposits unconformably overlie Permian limestone in the study area. The sandstones of the Paleocene Hangu Formation were followed by the deposition of the fully marine carbonates of the Lockhart Limestone, and, finally, by the Patala Formation [31,37]. The interplay of tectonic uplift, subaerial exposure, non-deposition, and the paleogeographic position of the Indian Plate around the equator [106,107] created suitable conditions for the development of bauxite and laterite in the central–eastern Salt Range (Figure 11a,b) during the Late Cretaceous to Early Paleocene.

Mineralogical and textural attributes suggest that the bauxites and laterites likely have a polygenetic origin: (1) a basal in situ ferrallitic weathering (lateritization) phase producing autochthonous-to-para-autochthonous material rich in kaolinite, and (2) overlying pisolitic conglomeratic layers deposited through reworking processes and resulting in deposits of allochthonous-to-para-allochthonous origin (Figure 11). L-1 represents predominantly autochthonous development. Field, SEM/EDS, and XRD analyses confirm that these layers originated as sandstone, limestone, or clay and subsequently underwent in situ ferrallitic weathering. The tectonic uplift and paleogeographic position of the Indian Plate during the Early Paleocene likely promoted lateritic (ferrallitic) weathering [37,112]. However, the ongoing Himalayan orogeny (India–Asia collisions) disrupted the long-term, stable surface conditions necessary for complete bauxitization, resulting in immature bauxite profiles with partial mineralogical transformation, shown by low gibbsite content and poorly developed textures, such as the absence of pisolitic structures. The occurrence of iron-rich nodules and aluminum hydroxide-encrusted clast fragments lacking sharp contact with the matrix suggests limited in situ weathering followed by, perhaps, repeated minor reworking [1] (Figure 5 and Figure 11c).

In contrast, the overlying L-2 exhibits varied ore textures, ranging from oolitic and pisolitic to conglomeratic. The presence of grains and pebbles of obviously older (texturally more “mature”) bauxite, along with angular hematite–goethite fragments among the ooids and pisoids, indicates that these bauxite horizons originated from the erosion and re-sedimentation of pre-existing bauxites or duricrusts [1,7]. The matrix is pelitomorphic to microgranular, containing microclastic and bauxitic pebbles surrounded by accretionary crusts. Most ooid and pisoid boundaries exhibit sharp contacts with the enclosing matrix (Figure 5c–e and Figure 11c), indicating a parautochthonous-to-allochthonous origin. In addition, the enrichment of ultra-stable heavy minerals indicates long-distance sediment transport and/or intense weathering, resulting in the selective removal of less stable (heavy) minerals [14,113]. L-3 reflects a change in depositional conditions: ferrallitic weathering remained active, as indicated by the presence of kaolinite, goethite, and anatase [114], but the process had either insufficient time or less favorable conditions to develop into proper bauxite ore (Figure 5f and Figure 11c).

6. Conclusions

• Field observations, mineralogical analysis, and geochemical data indicate that the bauxites of the Salt Range belong to the karst bauxite category. The ternary diagram of major oxides further classifies them as bauxite/bauxitic clay deposits.
• Hematite, goethite, and kaolinite are the primary constituents of L-1, with minor amounts of mica, quartz, boehmite, rutile, and anatase. L-2 contains boehmite, kaolinite, alunite/natroalunite, and zaherite, along with minor amounts of quartz, mica, rutile, and anatase, while L-3 predominantly comprises kaolinite and quartz, with rutile, goethite, and anatase as accessory minerals.
• Binary diagrams of geochemical proxies (Eu/Eu* vs. Sm/Nd and Eu/Eu* vs. ${\mathrm{TiO}}_2$/${\mathrm{Al}}_2$${\mathrm{O}}_3$) suggest a provenance from UCC and cratonic sandstones. Additionally, the Zr-Cr-Ga ternary diagram shows that all samples from ATB and AKB, and half of those from CMB, are derived from acidic and intermediate/argillaceous source rocks.
• Field studies, along with microscopic and SEM/EDS observations, reveal that the bauxite ore formed through multiple stages involving both diagenetic and late epigenetic processes. The L-1 facies reflects an autochthonous-to-para-autochthonous origin, while the L-2 facies are characterized by an allochthonous-to-para-allochthonous origin, pointing to an erosional event that followed a period of relative landscape stability, which previously resulted in the in situ ferrallitic weathering mentioned above.
• Further, geological evidence (including the geochemistry and mineralogy of the studied outcrops) indicates that these bauxite deposits are generally of sub-economic grade, except for a few particular layers within the conglomeratic facies.